Cruz Salmeron et al.
Progress in Earth and Planetary Science(2022) 9:44
https://doi.org/10.1186/s40645-022-00503-5

RESEARCH ARTICLE

Progress in Earth andPlanetary Science

Open Access

Characterization of water masses
around the southern Ryukyu Islands basedon isotopic compositions
Andros Daniel Cruz Salmeron1 , Hideko Takayanagi1 , Shigeyuki Wakaki2 , Tsuyoshi Ishikawa2 ,Toshihiro Miyajima3 , Hitomi Wakaki2 , Takuya Itaki4   and Yasufumi Iryu1*   

Abstract
We investigated the water-mass structure on the Okinawa Trough and Pacific sides of the southern Ryukyu Island Arc(Yonaguni, Iriomote, and Ishigaki subareas) using the Nd isotope composition (143Nd/144Nd ratios; expressed as εNdvalues) of benthic foraminiferal tests in surface sediments, which reflect bottom-water composition, along with hydrogen and oxygen isotope compositions (δD and δ18O values, respectively) and physical properties (temperature andsalinity) of seawater. The Okinawa Trough side has lower εNd values than the Pacific side due to continental/islandmaterial inputs characterized by relatively low εNd values. Moreover, within the Okinawa Trough, other processes control the Nd behavior of seawater and primarily affect the Yonaguni and Iriomote subareas, as follows. (1) Surface andsubsurface waters are influenced by Taiwanese river discharge combined with temporospatial variations in oceanographic conditions including Kuroshio Current meandering. (2) Intermediate water is characterized by low εNd values(down to − 8.2), possibly attributable to sediment plumes and turbiditic fluxes. (3) The εNd values of bottom water
indicate upwelling and vertical mixing, with composition therefore being similar to those of intermediate water. The
εNd profiles are better defined on the Pacific side. High εNd values occur in surface and subsurface (< 300 m depth,
potential density < 25.0 kg ­m−3) waters, and low values (down to − 7.0) occur in subsurface–core-intermediate water(400–600 m depth, 26–27 kg ­m−3). εNd values increase slightly to − 4.0 below 750 m depth and remain constantdown to about 2000 m depth, below which deep water shows a slight decrease in εNd values. Intermediate and bottom/deep waters are distinguished from upper layers by their lower δD and δ18O values.
Keywords: Neodymium isotopes, Hydrogen isotopes, Oxygen isotopes, Kuroshio Current, Okinawa Trough, Ryukyu
Islands, Northwestern Pacific*Correspondence: yasufumi.iryu.d8@tohoku.ac.jp1
Institute of Geology and Paleontology, Graduate School of Science, Tohoku
University, Sendai, Japan
Full list of author information is available at the end of the article

1 Introduction
The Kuroshio Current (KC) enters the East China Sea(ECS) through the strait between Taiwan and Yonaguni
Jima Island (TY), flows northeastward in the Okinawa
Trough (OT) along the Ryukyu Island Arc (RIA),
and bifurcates to the southwest of Yaku Shima Island(Fig. 1). The main current changes direction, veering
towards the Pacific in the Tokara Strait (TS) area, where
the Tokara Gap (TG) lies between Amami Oshima andYaku Shima islands. The KC is a warm current thattransports heat from the tropics to the Ryukyu Islands(the Ryukyus) and the southern half of mainland Japan.

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Cruz Salmeron et al. Progress in Earth and Planetary Science(2022) 9:44

It has enabled the development of reefs with highly
diversified hermatypic coral fauna throughout the
Ryukyus, despite their relatively high latitudes within
the reef province (Veron 1992) since the opening of
the OT and subsequent influx of the KC into the backarc basin (Iryu et al. 2006). Knowledge of the pathway,strength, and physicochemical properties of the KC is
critical for understanding modern oceanographic andclimatic conditions in East Asia, and a detailed study ofwater-mass characteristics around the RIA is expected
to contribute to this knowledge.
Neodymium (Nd) is a rare-earth element (REE) withseven naturally occurring isotopes: 142Nd (27.13% abundance), 143Nd (12.18%), 144Nd (23.8%), 145Nd (8.3%), 146Nd(17.19%), 148Nd (5.76%), and 150Nd (5.64%). Unlike thestable isotope ratios of light elements (e.g., hydrogen, carbon, and oxygen), 143Nd/144Nd ratios, expressed as εNdvalues (Sect. 3.2), are not affected by isotope fractionation through physical or biogeochemical process, biological cycles, ocean–atmosphere interactions, or seawatertemperature changes (Goldstein and Hemming 2003;
Lynch-Stieglitz 2003; Klevenz et al. 2008). Because Nd in
the oceans is supplied from continents and ocean floors
when a water mass is formed, it acquires the εNd valuesof continent- and ocean-floor-forming rocks. The Nd residence time in oceans (500–1000 yrs) is shorter than thetimescales of global oceanic circulation and homogenization (1500–2000 yrs; Tachikawa et al. 2003; Lacan andJeandel 2005). As a result, the εNd values of the oceansvary with depth and location (Piepgras and Jacobsen
1988; Goldstein and Hemming 2003; Lacan and Jeandel2005) and are stratified in the water column according to
the presence of different water masses (Amakawa et al.
2004b). Therefore, εNd values have been used to identify
water masses and reconstruct ocean circulation (Piepgras
and Jacobsen 1988; Amakawa et al. 2004b; Wu 2014; Wu
et al. 2015a).
Planktic foraminiferal tests record the εNd values ofambient seawater at various depths during their lifespan.
After planktic foraminifers die and their tests are transported to the sediment–water interface, the original εNdvalues are modified to reach isotopic equilibrium withbottom waters over at least 300 yrs (Palmer 1985; Roberts
et al. 2012). However, benthic foraminiferal tests comprise calcite precipitated in Nd isotope equilibrium with
ambient seawater during both their life and death stages.
Because benthic foraminifers also have wide temporospatial distributions, the εNd values of benthic foraminiferal
tests have been used to detect variations in the εNd values of bottom water due to water-mass structures (Murray 2013). There is no significant difference in εNd values
between a sample of a single benthic foraminiferal species and a sample of multiple foraminiferal species (Haley

Page 2 of 20

et al. 2005; Roberts et al. 2012; Wu et al. 2015a); therefore, when single-species samples are limited or unavailable, the εNd values of multi-species samples can be used
as a reliable proxy for identifying a water mass and its
temporospatial variation (e.g., Palmer 1985; Lynch-Stieglitz and Marchitto 2014).
The ECS is influenced by multiple sources of Nd. The
primary source is the Philippine Sea (PS), where εNdvalues range from − 4.4 in deep water to − 2.0 in surface–subsurface water (Wu 2014; Behrens et al. 2018).
Although in lower proportions, some influence of theSouth China Sea (SCS) is known to reach the ECS
through the TY and the Kerama Gap (KG; Chen 2005;
Nakamura et al. 2013). The εNd values in the northeastern SCS vary from − 4.5 in deep water to − 3.6 in subsurface water, with a surface-water εNd value of − 4.5. Manyrivers in China flow into the ECS, with the most significant influence being from the Changjiang (εNd =  − 11.6),
Yellow (− 13.9 to − 9.6), and Yangtze (− 14.3 to − 8.7) rivers (Meng et al. 2008; Dou et al. 2012). The OT is also
influenced by rivers in Taiwan (εNd =  − 11.1; Lan et al.
2002).
There have been few Nd isotope studies focusing
on the OT, and even fewer on the KC. Amakawa et al.
(2000) reported εNd values of Kuroshio Surface Water(KSW) at < 5 m depth southeast of Taiwan (εNd =  − 4.8
at station PA-S-18; 21.83°N, 121.11°E; Fig. 1) and in thewestern–central OT (− 4.9 at station PA-S-19; 26.30°N,
125.54°E; Fig. 1), with these values being lower than thosemeasured to the east of Luzon Strait in the PS (− 4.2 atstation PA-S-1; 19.54°N, 131.27°E; Fig. 1). A Nd isotope
profile at station LM-6/11 (34.16°N, 142°E; Fig. 1) exhibits lower values than those measured to the east of Taiwan (Amakawa et al. 2004b), with εNd values of − 5.6 at
5 m depth, − 8.7 at 251 m, − 6.0 at 503 m, and − 5.9 at
1989 m. This SW–NE decrease in εNd values in the KC
has been attributed to riverine sediment inputs fromthe ECS (Amakawa et al. 2004b; Behrens et al. 2018)
and transportation of less radiogenic continental Nd(relatively low εNd values) into the central Pacific by
the KC (Amakawa et al. 2004a). Dou et al. (2012) identified a post-glacial increase in εNd values during the lastdeglaciation (from − 11.4 at 14 ka to − 10.7 at 7 ka) in
sediments collected at 766 m water depth on the slopeof the ECS shelf, and they interpreted this as reflecting
a reduced influence of riverine input from China due tostrengthening of the KC. However, no studies have determined the Nd isotope composition in the KC around the
RIA at high spatial resolution and over a wide range of
depths, with most previous studies focusing on the KC
before entry or after exit of the OT.
Stable hydrogen and oxygen isotope compositions,
D/H and 18O/16O expressed as δD and δ18O values,

Cruz Salmeron et al. Progress in Earth and Planetary Science(2022) 9:44

Page 3 of 20

ECS
in

pu

t

KC

TG

SC
S

wa
te
r

Tai
wa
n

S
RC yu

Ry

TY

uk

Ar
c

d
N

P
W

PS
DW
UC
KC

in

pu

t

DW
UC

LS

SCS

a
wKG
a
n

gh

Is
l

O

ki

ou
Tr

an

KC

Co
n

tin

en

ta
l

China

P
N DW
PI
W

Japan

Pacific Ocean

NEC

Fig. 1 Surface and subsurface current systems in the northwestern Pacific Ocean. Red points represent stations from which Nd isotope data havebeen reported in previous studies: LM-6/11 (Amakawa et al. 2004b); PA-S-1, PA-S-18, and PA-S-19 (Amakawa et al. 2000); 2, 3, 4, and 5 (Behrens et al.
2018); LS C3a, P2, SCS C5, and WPS N18-05 (Wu 2014). Abbreviations: ECS, East China Sea; KC, Kuroshio Current; KG, Kerama Gap; LS, Luzon Strait;
NEC, North Equatorial Current; NPIW, North Pacific Intermediate Water; PDW, Pacific Deep Water; PS, Philippine Sea; RCS, Ryukyu Current System;
SCS, South China Sea; TY, Strait between Taiwan and Yonaguni Jima Island; TG, Tokara Gap; UCDW, Upper Circumpolar Deep Water. White points
represent stations for which δD values were reported by Horibe and Ogura (1968) (Additional file 5: Fig. S3). The horizontal yellow line in the Luzon
Strait indicates a transect studied by Mensah et al. (2014)

respectively (Sect. 3.2), provide crucial oceanographic
information regarding thermohaline circulation and theformation, mixing, and ventilation of water masses. δDand δ18O values vary spatially in the oceans depending on
advection processes and the mixing of water masses withdifferent isotopic compositions (Xu et al. 2012; Rohling2013). Transition phases of the hydrological cycle, turbulence fluxes, and Rayleigh distillation affect the δD andδ18O values of surface water, which change progressively
from low to high latitudes. For example, heavier isotopic
compositions occur in areas of high evaporation near theEquator, whereas lighter compositions are observed near
the poles (Oka and Kawabe 1998; Schmidt et al. 1999;
Rohling 2013). The relationships of temperature andsalinity to δD and δ18O values have been broadly used to
identify water masses and ocean dynamics (e.g., Horibeand Ogura 1968; Schmidt et al. 1999; Bigg and Rohling,
2000; LeGrande and Schmidt 2006; Sengupta et al. 2013;
Wu et al. 2021). Global datasets indicate lower δ18O values near coastal areas such as the ECS and the PS, whichhave δ18O values of –0.1‰ and + 0.1‰, respectively (Biggand Rohling 2000; LeGrande and Schmidt 2006).
Evaporation and precipitation control salinity–δD andsalinity–δ18O relationships of shallow water distal to continental inputs. When evaporation exceeds precipitationand freshwater discharge, δ18O and δD values and salinity
increase. Conversely, when precipitation and freshwater
discharge exceed evaporation, δ18O and δD values andsalinity decrease (Rohling 2013; Sengupta et al. 2013).
Furthermore, physical oceanographic processes suchas sinking, advection, and upwelling may cause different water masses to have distinct slopes and intercepts

Cruz Salmeron et al. Progress in Earth and Planetary Science(2022) 9:44

on salinity–δD and salinity–δ18O cross-plots (Sengupta
et al. 2013).
The KC is a source of warm oligotrophic water, so its
influx into the OT has been important for the evolutionof oceanographic conditions and climate states in Far
East Asia. However, past flow paths of the KC and thetiming of KC entrance to the OT are poorly constrained.
Existing proxies, such as δ13C and δ18O values of microfossils, numerical models, and sediment provenanceshow some discrepancies (e.g., Ujiié et al. 2003; Kao et al.
2006; Dou et al. 2010, 2012; Lee et al. 2013). This study
aimed to establish a new proxy of Nd isotope composition that may be used to reconstruct the past water-massstructures of the OT and around the RIA. From a physical oceanography perspective, the OT is a challenging
study area due to its complex bathymetry, seasonal-tointerannual variability in seawater physical properties,
occurrence of several water masses, and riverine inputs.
Therefore, Nd isotopes are expected to provide vital
information on the current system and water-mass structure within the OT.
In this study, εNd values were determined by analysis
of a small number of modern benthic foraminifers, andthe values are directly compared with εNd data from previous studies. This is the first study to delineate variationsin εNd values with wide spatial coverage in the KC andrelated water masses on the OT and Pacific sides of thesouthern RIA at high spatial resolution, based on εNd,
δD, and δ18O values and physical properties of seawater(e.g., temperature and salinity). The study area is locatednear the KC entrance to the OT, so it is an ideal site forcharacterization of the Nd isotope signature of the KC
immediately after it interacts with terrigenous materials
from the Eurasian continent and Taiwan.

2 Study area
2.1 Geological settings

The Ryukyus include more than 200 islands and islets
to the southwest of mainland Japan, extending from
Tanegashima Island (30.73°N, 131°E) in the northeast to
Yonaguni Jima Island (24.45°N, 123°E) in the southwest(Fig. 1). The islands are arranged in an arc known as the
RIA, an active island arc related to subduction of thePhilippine Sea Plate along the Ryukyu Trench beneaththe Eurasian Plate. The RIA is bounded by the ECS to
the northwest and the Pacific Ocean to the southeast
and is geographically divided into three regions (North,Central, and South Ryukyus; Iryu et al. 2006; Arai et al.
2014) by the TG and KG (Fig. 1). The OT is a back-arc
basin separating the RIA from the ECS shelf. The study
area is located in the western part of the South Ryukyus.
The study area on the OT side is divided into three subareas (Fig. 2): Yonaguni (Yonaguni OT), Iriomote (Iriomote

Page 4 of 20

OT), and Ishigaki (Ishigaki OT). On the Pacific side, it
is also divided into three subareas: Yonaguni (Yonaguni
Pacific), Iriomote (Iriomote Pacific), and Ishigaki (Ishigaki Pacific).
The OT has a NE–SW length of ~ 1000 km and a NW–
SE width of 200 km. The bathymetry of the trough ranges
from shallower depths of ~ 600 m in the north to maximum depths of ~ 2100 m in the Yaeyama and Miyako rifts
in the south. Three channels connect the OT with the PS;
i.e., the TY, KG, and TG, with maximum water depthsof 1700, 2100, and 1500 m and sill depths of 775, ~ 1100,
and 700 m, respectively (Arai et al. 2013; Nakamura et al.
2013; Nishizawa et al. 2019).
The opening of the OT and subsequent influx of the KC
into the back-arc basin was a key event in the Cenozoic
environmental and oceanographic evolution of Far East
Asia including the Ryukyus, with the KC being a sourceof warm oligotrophic water. The initial formation of coral
reefs in the Ryukyus is considered to record the influx of
the KC into the ECS (i.e., the back-arc side of the RIA).
This changed the oceanographic conditions from a “mudsea,” represented by siliciclastic deposits of the upper
Miocene–lower Pleistocene Shimajiri Group, to a “coralsea,” represented by carbonates of the Pleistocene Ryukyu
Group that formed in reefs and associated fore-reefs andisland shelves (Iryu et al. 2006; Gallagher et al. 2015; Imaiet al. 2017).
2.2 Oceanographic settings

Here we introduce general aspects of the ocean circulation and water masses in the study area, referring tooceanographic data for the three subareas on the OT
and Pacific sides of the RIA (Additional file 1: Tables S1,
Additional file 2: Table S2; Fig. 2, Additional file 3: Fig.
S1) provided by the Japan Oceanographic Data Center(Japan Oceanographic Data Center, 2021). The JODC
temperature and salinity data used in this study werecollected from 1 January 1990 to 31 December 2019.
Long-term (30 yr) averaging of the data has eliminatedshort-term variations in temperature and salinity, suchas those caused by seasonality, the El Niño–Southern
Oscillation, and variability in the KC flow path (Nitani
1975). Therefore, the JODC data are more suitable forcharacterization of the water masses investigated here
than onboard data obtained during the GK19 cruise(Additional file 1: Tables S1, Additional file 2: Table S2).
We identified water masses following water-mass classifications of previous studies (Nakamura 1996; Wang andChen 1998; Qu et al. 1999; Ichikawa and Chaen 2000;
Amakawa et al. 2004b; Chen 2005; Talley 2008; Kawabe
and Fujio 2010; Talley et al. 2011; Amakawa et al. 2013;
Qi et al. 2014; Wu 2014; Yang et al. 2015; Behrens et al.
2018; Zuo et al. 2019). Using the JODC salinity andCruz Salmeron et al. Progress in Earth and Planetary Science(2022) 9:44

temperature data, we calculated potential temperaturesand densities and then correlated our results using previous data for the OT and PS (Tables 1 and 2).
2.2.1 Okinawa trough side of the Ryukyu Island Arc

The North Equatorial Current (NEC) bifurcates into thenorthward-flowing KC and southward-flowing Mindanao
Current to the east of the Philippine coast. The KC enters
the OT, flows northeastward along the RIA, and drifts
out to the Pacific through the TS. The core of the KC has
an average thickness of 500–700 m but reaches depthsof up to 1350 m in some areas (Chen 2005; Kubota et al.
2017). Based on physical and chemical properties (e.g.,temperature, salinity, nutrients, and pH), three water
masses have been identified within the OT in the verticalprofile of the KC path: KSW (< 150 m depth), Kuroshio
Tropical Water (KTW; 150–300 m depth), and Kuroshio
Intermediate Water (KIW; 500–700 m depth; Chen et al.
1995; Wang and Chen 1998; Chen 2005; Mensah et al.

Page 5 of 202014). Distinctive KC temperature and salinity fronts
exist between the continental shelf region and its offshore region in the OT, with salinity being lower in theformer due to freshwater discharge from rivers in China(Oka and Kawabe 1998).
The JODC temperature and salinity data indicate thepresence of KSW above ~ 100 m depth, with salinity of
34.47–34.78 and potential density (σθ) of < 23.0 kg ­m−3.
KTW corresponds to the maximum salinity layer at
100–200 m depth, with salinity of 34.81–34.88 and σθ
of 23.5–24.6 kg ­m−3. Western North Pacific Central
Water (WNPCW) is present below the KTW at 250–
500 m depth, with salinity of 34.32–34.82 and σθ of
25.0–26.4 kg ­m−3 (Table 1, Fig. 4).
Intermediate water resulting from horizontal advection and mixing between North Pacific Intermediate Water (NPIW) and South China Sea Intermediate
Water (SCSIW; Nakamura et al. 2013) enters the OT
via the KG and TY (Chen 2005; Nakamura et al. 2013;

Okinawa Trough

wate
SCS

TY

RCS

r

TaiwanLanyang
River

KC

Yonaguni

Iriomote

Ishigaki

Pacific Ocean

Fig. 2 Sampling sites in the study area. Circles and triangles represent sites on the Okinawa Trough and Pacific sides of the southern Ryukyu Island
Arc, respectively. Green, blue, and orange symbols indicate sites in the Yonaguni, Iriomote, and Ishigaki subareas, respectively. Circles and triangles
with red borders indicate sites where sediment samples were collected for Nd isotope analysis. For abbreviations, see Fig. 1

Cruz Salmeron et al. Progress in Earth and Planetary Science(2022) 9:44

Na et al. 2014). As a result, the salinity of KIW falls
to levels between those of NPIW and SCSIW. KIW
exists at 500–755 m depth in the OT side, characterized by its low salinity of 34.29–34.36 and σθ of 26.5–
27.0 kg ­m−3. The intermediate water is more salineby ~ 0.11 on the OT side of the RIA than on the Pacificside, indicating mixing with more saline water within
the OT, as for SCSIW (Fig. 3). In the southern OT, themixing proportions of SCSIW and NPIW are 55% and45%, respectively (Nakamura et al. 2013). A salinity
front separating western high-salinity water from eastern low-salinity water has been detected to the east of
Taiwan at 21.45°N near 122.5°E (Chen 2005), suggesting that SCSIW encounters NPIW in this area (Fig. 1).
Bottom waters in the OT have salinity of 34.63 and σθ
of > 27.5 kg m
­ −3 (Table 1; Fig. 4).
2.2.2 Pacific side of the RIA

North Pacific Subtropical Gyre water flows northeastward across the PS on the Pacific side of the RIA, and thesubsurface layer has been identified as the Ryukyu Current System (Nakamura et al. 2007; Thoppil et al. 2016).
The JODC data indicate that the salinity and σθ of surface
water (< 83 m depth) are 34.31–34.74 and < 23.00 kg ­m−3
on the Pacific side of the RIA, respectively. However,these values vary seasonally, with salinity being lower insummer and higher in winter (Rudnick et al. 2011). Moreover, the salinity of surface–subsurface water increases inan E–W direction (Chen 2005; Mensah et al. 2014). Salinity maxima correspond to NPTW at 99–216 m depth in
the study area (34.67–34.90; σθ = 23.5–24.7 kg ­m−3).
Central Water, formed in the subtropics and Equatorial Pacific, overlies North Pacific Intermediate Water

Page 6 of 20(NPIW) and is characterized by a wide range of potential
temperature (θ, °C) and salinity. Central Water spreads
over the thermocline through advection and diffusion
processes over the isopyncals after subduction from thesurface mixed layer. For North Pacific Central Water,
upper and lower limits correspond to North Pacific
Subtropical Mode Water (σθ = 25.5 kg ­m−3) and North
Pacific Central Mode Water (σθ = 26.3 kg ­m−3), respectively (Nakamura, 1996; Talley et al., 2011). Here we referto this water mass as being WNPCW, identified as having salinity of 34.37–34.82 and σθ of 24.9–26.0 kg ­m−3 at
222–476 m depth (Fig. 4).
NPIW has been identified as the minimum-salinity
layer in the western PS (Wang and Chen 1998; Qu et al.
1999; Chen 2005; Behrens et al. 2018; Table 2). In the
study area, NPIW has salinity of 34.18–34.39 and σθ of
26.3–27.2 kg ­m−3.
In the Pacific Ocean, the southern-sourced Lower Circumpolar Deep Water and Upper Circumpolar Deep
Water (UCDW) flow in a S–N direction. UCDW upwells
at high latitudes in the North Pacific and flows southward
as Pacific Deep Water (PDW; Talley et al. 2011). In thePS, PDW can be traced at > 2000 m depth with salinityof 34.6–34.7, θ of 1.6 °C–2.1 °C, and σθ of > 27.6 kg ­m−3
(Amakawa et al. 2004b, 2013; Talley 2008; Wu et al.
2015a; Table 2). PDW flows above UCDW to the eastof Luzon Strait in the western PS (Wu et al, 2015b;
Wan et al. 2018), whereas it is at 1907–2620 m depth
in the study area, where salinity is 34.59–34.64 andσθ > 27.7 kg ­m−3.
UCDW flows northward in the PS at 3000–3500 m
depth (Kawabe and Fujio 2010), encountering the southward-flowing PDW with θ = 1.55 °C and salinity 34.64

Table 1 Characterization of the water masses in the southern Ryukyu Island Arc. *Only one data available
Depth

T (°C)

Salinityσθ (kg m­ −3)

δD (‰,
average)δ18O (‰,
average)

εNd

Kuroshio Surface Water< 100

23.4–28.2

34.47–34.78< 23.04.5*

0.5*

Kuroshio Tropical Water

100–200

18.6–24.1

34.81–34.88

23.5–24.6

3.6

0.4

− 2.2

Western North Pacific Central Water

250–500

10.0–18.6

34.32–34.82

25.0–26.4

3.1

0.4

Kuroshio Intermediate Water

500–755

5.8–9.0

34.29–34.36

26.5–27.0

1.0

0.2

Deeper Kuroshio Intermediate Water

755–1900

3.6–5.7

34.36–34.48

27.0–27.5

0.4

0.0

Deep water

1900–1938

2.0

34.63

> 27.5––

Surface water< 83

24.5–28.6

34.31–34.74< 23.03.3

0.4

North Pacific Tropical Water

99–21619.3–24.6

34.67–34.90

23.5–24.7

3.4

0.4

Western North Pacific Central Water

222–476

11.2–18.9

34.37–34.82

24.9–26.0

3.1

0.4

North Pacific Intermediate Water

478–863

4.9–11.4

34.18–34.39

26.3–27.2

1.3

0.1

Deeper North Pacific Intermediate Water

864–1851

2.2–4.9

34.40–34.61

27.2–27.6

0.5

0.1

Pacific Deep Water

1907–2620

1.7–2.1

34.59–34.64> 27.70.7

0.0

Okinawa Trough side− 5.3 to − 5.2− 6.2 to − 2.8− 5.1 to − 4.8− 8.2 to − 4.9–

Pacific side− 1.7 to − 1.2− 4.1 to − 1.1− 5.2 to − 3.4− 7.0 to − 3.8
− 4.0

− 5.3

Cruz Salmeron et al. Progress in Earth and Planetary Science(2022) 9:44

Page 7 of 20

Table 2 Water masses previously reported from the Okinawa Trough and the Pacific Ocean
Depth

Pot. Temp. (°C) Salinityσθ (kg/m3) εNd

Surface Water

0–50

26.3–28.71> 34.78> 22.48

North Pacific Tropical Water

61–25018.67–24.50

34.75–35.25 24.0–24.5

Western North Pacific Central Water

250–400

8.5–17.5

34.10–34.80 25.3–26.5

North Pacific Intermediate Water

500–700

5.13–8.22

34.15–34.4

26.5–27.2− 3.3 to − 2.6 Wang and Chen (1998), Qu et al.
(1999), Chen (2005), Behrens et al.
(2018)

Pacific Deep Water> 20001.55–2.1

34.6–34.7

27.6− 6.0 to − 2.8 Amakawa et al. (2004b), Talley (2008),
Amakawa et al. (2013), Behrens et al.
(2018)

Upper Circumpolar Deep Water,
southern Pacific

3000–4000

1.6–2.1

34.6–34.7

27.55–28− 6.5 to − 4.6 Talley (2008), Kawabe and Fujio (2010),
Amakawa et al. 2013)

South China Sea Surface Water

0–10

24.9–26.3

33.9–34.4

27.55− 4.5 to − 2.9 Wu (2014), Japan Oceanographic Data
Center (2021)

South China Sea Intermediate Water

500–600

7.43–8.38

34.38–34.43 26.25–27.0 − 4.1 to − 3.6 Wu (2014), Japan Oceanographic Data
Center (2021)

Kuroshio Surface Water

0–75

14–30.2

34.2–35.1

22.85

–

Ichikawa and Chaen (2000), Qi et al.
(2014), Zuo et al. (2019)

Kuroshio Tropical Water

100–300

17–25

34.85–35.1

23.5–24–

Chen (2005), Ichikawa and Chaen(2000),Yang et al. (2015)

Kuroshio Intermediate Water

400–800

5–13

34.1–34.4

26.5–26.77 –Chen, (2005), Ichikawa and Chaen(2000), Yang et al. (2015), Zuo et al.
(2019)

Water masses

References

Pacific Ocean− 3.5 to − 2.0 Amakawa et al. (2004b), Wu (2014),
Behrens et al. (2018)− 3.1 to − 2.1 Behrens et al. (2018)− 3.7 to − 3.2 Nakamura (1996), Talley et al. (2011),
Behrens et al. (2018)

South China

Okinawa Trough(station LM-6/11; Amakawa et al. 2004b) at 27°–30°N(Kawabe et al. 2009). However, it is difficult to separate
them within the PS at 1000–2000 m depth because a portion of UCDW upwells (Kawabe and Fujio 2010) and isvertically mixed with PDW. Some authors therefore referto them as UCDW/PDW (e.g., Behrens et al. 2018).

3 Materials and methods
3.1 Sampling and sample processing

A total of 189 bottom-water and sediment samples werecollected using a K-grab sampler from the area between
Ishigaki Jima and Yonaguni Jima islands over a depth
range of 68–2620 m during the GK19 cruise of R/VKaiyo-maru No. 1 undertaken during June 23–July 22,
2019, by the Geological Survey of Japan of the National
Institute of Advanced Industrial Science and Technology (AIST/GSJ; Additional file 1: Tables S1, Additional
file 2: Table S2; Fig. 2). Bottom-water samples were collected from ~ 7 m above the seafloor using a Niskin sampler attached to a K-grab sampler. Some physical andchemical properties were measured using a conductivity–temperature–depth (CTD) profiler (CTD 90 M; Sea& Sun Technology GmbH, Germany) attached to thegrab sampler. Temperature, salinity, turbidity, pH, anddissolved-oxygen data are presented in Additional file 1:
Tables S1 and Additional file 2: Table S2. Surface sediment samples (5–10 ­cm3; 0–2 cm depth) collected from42 sites were washed using ultra-pure water and a 63 µm
mesh sieve. Particles of > 63 µm size were air-dried at
40 °C for 12 h and sieved using 125 and 250 µm sieves.
Samples were divided into three size fractions: 63–125,
125–250, and > 250 µm. Benthic foraminiferal tests (50–70) were handpicked from the > 250 µm fraction under abinocular microscope (Stemi 2000-C; Zeiss, Germany).
Although the ages of the samples were not determined,
the samples were obtained from 0–2 cm below the seafloor, and previous studies have dated such seafloor sediment at < 1 ka near the study site (Chang et al. 2008; Douet al. 2012). Figures were processed using QGIS 2.18.19,
Ocean Data View (Schlitzer, 2018), and Inkscape version
0.92 (https://​inksc​ape.​org/) software. Gridded bathymetric data were taken from the GEBCO Compilation Group(2020), from which isobaths were obtained every 200 m
from the digital elevation model using QGIS 2.18.19.
Geographical names follow the Gazetteer of Japan 2007
(Geographical Survey Institute of Japan and Japan Coast
Guard 2007).

Cruz Salmeron et al. Progress in Earth and Planetary Science(2022) 9:44

Page 8 of 20

Fig. 3 Depth profiles of salinity, temperature, and hydrogen (δD), oxygen (δ18O) and neodymium (εNd) isotope compositions for the Okinawa
Trough (upper panels) and Pacific (lower panels) sides of the southern Ryukyu Island Arc. The salinity and temperature profiles are based on thedata collected from 1 January 1990 to 31 December 2019 (Japan Oceanographic Data Center 2021). Wine-colored symbols in the salinity profile of
the Okinawa Trough side indicate salinity measured in the South China Sea water (Japan Oceanographic Data Center 2021). Solid lines in the δDand δ18O profiles indicate mean δD and δ18O values in individual water masses. Dotted lines indicate mean ± 2σ ranges. Horizontal bars in the Nd
isotope profiles represent analytical errors (2SD). Okinawa Trough side (upper panels): KSW, Kuroshio Surface Water; KTW, Kuroshio Tropical Water;
WNPCW, Western North Pacific Central Water; KIW, Kuroshio Intermediate Water; BW, Bottom water. Pacific side (lower panels): SW, surface water;
NPTW, North Pacific Tropical Water; WNPCW, Western North Pacific Central Water; NPIW, North Pacific Intermediate Water; PDW, Pacific Deep Water

Fig. 4 Potential temperature–salinity plots for the Okinawa Trough (left panel) and Pacific (right panel) sides of the southern Ryukyu Island Arc.
Data from the Japan Oceanographic Data Center (2021)

Cruz Salmeron et al. Progress in Earth and Planetary Science(2022) 9:44

3.2 Sediment sample preparation and Nd isotope analyses

Forty-two sediment samples were selected for Nd isotope analysis. The cleaning of bulk benthic foraminifers
followed Lea et al. (2000) and Sagawa et al. (2005). Bulk
benthic foraminiferal samples were collected in a 1.5 mL
micro-centrifuge tube and decontaminated using a multistep cleaning procedure, including clay removal using
ultrapure water and methanol, reductive cleaning to
remove oxidized metal coatings using hydrazine/ammonium citrate, and oxidative cleaning to remove organic
matter using ­H2O2 buffered by NaOH. The cleaning wasundertaken at the Institute of Geology and Paleontology,
Graduate School of Science, Tohoku University, Japan.
Nd isotope (143Nd/144Nd) analysis of sub-nanogram–nanogram Nd samples followed Wakaki and Ishikawa(2018). The Nd in bulk benthic foraminiferal samples
was chemically separated and purified using a three-step
extraction and ion-exchange chromatographic procedure. First, Nd was separated from major elements andother REEs using TRU and Ln resins (Eichrom, USA) following a procedure modified from that of Pin and Santos(1997). Sample solutions were first loaded onto Tefloncolumns filled with 0.2 mL of TRU resin. Major elements
were eluted with 4 mL 3 M ­HNO3, before light REEs
(LREEs) were eluted with 2 mL ­H2O. Eluted LREEs were
loaded directly onto a second Teflon column filled with
0.5 mL Ln resin. La, Ce, and Pr were eluted with 3.5 mL
0.2 M HCl, and the Nd fraction with 3 mL 0.2 M HCl.
Finally, the separated Nd fraction was loaded onto a third
Teflon column filled with 0.1 mL cation-exchange resin(AG50W-X12; Bio-Rad Laboratories, Inc., Japan) for Nd
purification. Impurities, remaining cations, and organic
materials that may affect Nd ionization during thermal
ionization mass spectrometry (TIMS) were eluted with
1.5 mL 2.5 M HCl, and the purified Nd fraction was collected in 2 mL 6 M HCl and evaporated to dryness.
The purified Nd was dissolved in 2.5 µL 3 M H­ NO3 andloaded onto degassed Re double filaments with 0.8 µL
0.3 M ­H3PO4. Nd isotope ratios were determined using
TIMS (Triton, Thermo-Finnigan, USA) at the Kochi
Core Center, Japan, using the total evaporation normalization (TEN)–TIMS technique, with N­ d+ ions used toanalyze the Nd isotope composition of a small amountof sample (Wakaki and Ishikawa 2018). The 143Nd/144Nd
ratio was corrected by internal normalization using the
exponential law with an 146Nd/144Nd ratio of 0.7219(O’Nions et al. 1977) to correct for mass discrimination;
143
Nd/144Nd ratios were corrected for a small systematic
bias based on the recommended 143Nd/144Nd value ofstandard JNdi–1 (0.512115; Tanaka et al. 2000). The final
analytical error for the normalized 143Nd/144Nd ratio,
expressed as 2σ (Additional file 1: Tables S1, Additional
file 2: Table S2), was estimated by error propagation ofPage 9 of 20

the counting errors for 143Nd, 144Nd, and 146Nd throughthe normalization equation. Results are expressed in εNdepsilon Nd units.

εNd =[(143 Nd/144 Ndsample )/(143 Nd/144 Nd)CHUR − 1] × 104 ,(1)

where (143Nd/144Nd)CHUR = 0.512638 (Jacobsen and Wasserburg 1980).
3.3 Water sample preparation and hydrogen and oxygenisotope analyses

Bottom-water samples for H and O isotope analyses werecollected at 189 sites at depths of 68–2620 m during the
GK19 cruise. Samples were filtered with a 0.45 µm membrane filter and collected in 30 mL glass vials, which weresealed immediately after collection and stored in a cool,dark place.
The hydrogen (D/H) and oxygen (18O/16O) isotope
ratios of the water samples were determined using cavity ring-down spectroscopy (L2120-I Analyzer; Picarro,
USA) at the Atmosphere and Ocean Research Institute,
The University of Tokyo, Japan, without chemical preprocessing. Measured H and O isotope ratios were calibrated against an in-house standard and converted into
the conventional Vienna standard mean ocean water(V-SMOW) scale (per mil; ‰) based on the following
equations:δD = [(H/Dsample )/(H/D)standard − 1] × 103(2)

δ 18 O = [(18 O/16 Osample )/(18 O/16 O)standard − 1] × 103 .

(3)
External precision for H–O isotope analyses, based on
replicate measurements (n = 38) of the in-house standard, was ± 0.61‰ for δD and ± 0.10‰ for δ18O.
Temperature and salinity were measured using the
CTD profiler at the bottom-water sampling sites; thepH of the bottom-water samples was measured using acompact pH meter (LAQUAtwin-pH-33; Horiba Scientific, Japan) immediately after samples were collected.
The external precision (1σ) of pH values, based on replicate measurements of a standard solution (pH = 7.41),
was ± 0.02.

4 Results
4.1 Neodymium isotopes

The εNd values ranged from − 8.2 to − 2.2 on the OT sideof the RIA (Fig. 3). Comparison of εNd profiles in thethree subareas shows (1) an eastward increase at < 300 m
depth in the Iriomote and Ishigaki OT; (2) the occurrenceof a water mass with a wide range of εNd values (− 5.9
to − 2.8) at 300–600 m depth; and (3) relatively constantCruz Salmeron et al. Progress in Earth and Planetary Science(2022) 9:44

εNd values (− 5.9 to − 4.8) at > 600 m depth, with two
highly negative values of − 7.6 and − 8.2 being recordedin the Yonaguni OT at depths of 848 and 1026 m,
respectively.
The εNd values ranged from − 7.0 to − 1.1 on the Pacificside of the RIA (Fig. 3). The Nd isotope profiles of thethree subareas on the Pacific side had similar values andtrends from surface to deeper waters, with εNd valuesof − 2.8 to − 1.1 at < 157 m depth (with an outlier of − 4.1
at 125 m in the Iriomote Pacific). At > 222 m depth, εNdvalues decreased to − 5.5 at 656 m depth, with an outlier of − 7.0 at 497 m at Yonaguni Pacific, and increasedto − 4.0 at 1239 m depth (Fig. 3). A value of − 5.3 wasrecorded at 2264 m depth at the Iriomote Pacific site.
4.2 Hydrogen isotopes

On the OT side of the RIA, seawater δD values had similar depth trends for the three subareas (Fig. 3), except at
250–500 m depth in the Ishigaki OT where higher values were recorded. δD values were higher at 117–388 m
depth (up to 4.9‰), decreasing steadily to a minimum
value of − 1.0‰ at 1093 m depth. Although average δDvalues were almost constant at greater depths, an outlierof > 1.3‰ was recorded at 1800 m depth.
The three subareas on the Pacific side also displayed similar trends, with values of up to 5.3‰ above
380 m depth. δD values decreased to their lowest valueof − 0.7‰ at 976 m depth, before gradually increasing to
1.8‰ at 2128 m depth. A value of − 0.8‰ was recorded at
2620 m depth (Fig. 3).
4.3 Oxygen isotopes

On the OT side of the RIA, seawater δ18O values had
similar depth trends in the three subareas (Fig. 3), withhigher values of up to 0.7‰ being recorded above 298 m
depth. Below this depth, values decreased to a minimum of − 0.2‰ at 760 m depth, remaining roughly constant at ~ 0.0‰ at greater depths with an outlier of 0.3‰
at 1800 m. The δ18O profiles on the Pacific side also
had similar trends in the three subareas, with δ18O values being higher above 285 m depth (up to 0.7‰) anddecreasing to − 0.2‰ at 726–976 m depth before increasing to 2.0‰ at 1850 m depth (Fig. 3).

5 Discussion
5.1 Nd isotopes
5.1.1 Surface–subsurface waters

The εNd values at < 157 m depth (σθ < 23.5 kg ­m−3)
were − 5.3 and − 5.2 in the Iriomote OT and − 2.2 in the
Ishigaki OT; no data were available for the Yonaguni OT.
Similarly, low εNd values (− 4.3 to − 6.2) were recorded

Page 10 of 20

at 247–500 m depth (σθ = 25.0–26.4 kg ­m−3) in the threesubareas on the OT side of the RIA. Surface-water (5 m
depth) εNd values of − 4.8 and − 4.9 have been reported
previously southeast of Taiwan (PA-S-18) and west of
Okinawa Jima (PA-S-19) (Amakawa et al. 2000; Fig. 1).
The εNd values of the KC decrease as it flows to higher
latitudes in the OT and then to the southeast of Honshu,
as shown by low εNd values of − 7.3 to − 3.9 (< 200 m
depth, station 2; Behrens et al. 2018) and − 8.7 to − 5.6(< 455 m depth, station LM-6/11; Amakawa et al. 2004b;
Figs. 1 and 5).
When Northwest Pacific waters flow near Taiwan orinto the ECS, they are exposed to the influence of sediment and water with low εNd values derived from theEurasian continent and Taiwan, as indicated by εNd values as low as − 6.2 at 0–500 m depth (σθ < 26.4 kg ­m−3) on
the OT side of the RIA. Possible low-εNd sources include
continental rocks of the Eurasian continent (China) andTaiwan, which are enriched in REEs and have strongly
negative εNd values of − 15.2 to − 8.2 (Lan et al. 2002; Liet al. 2013; He et al. 2015). Sediments derived from theserocks are transported to the ocean by numerous rivers,
including the Changjiang (εNd =  − 11.6; Dou et al. 2012),
Yellow (− 13.9 to − 9.6; Meng et al. 2008), and Yangtze(− 14.3 to − 8.7; Meng et al. 2008) rivers on the Eurasiancontinent and Taiwan rivers (− 11.1; Lan et al. 2002). In
contrast, surface water (εNd =  − 1.7 to − 1.2) and NPTW(− 3.4 to − 1.1) on the Pacific side of the study area retain
the Nd isotope signatures of those from the PS, for whichsimilar εNd values (Fig. 5, Table 2) have been reported;e.g., − 3.2 to − 2.0 (Behrens et al. 2018; Wu 2014).
On the OT side of the RIA, WNPCW is characterized by εNd values of − 6.2 to − 2.8 at 250–500 m depth(σθ = 25.0–26.4 kg ­m−3). On the Pacific side, this watermass has εNd values of − 5.2 to − 3.4 at 222–476 m
depth (σθ = 24.9–26.0 kg ­m−3; Fig. 3). In the western PS,
North Pacific Subtropical Mode Water at 200–300 m
depth (σθ = 25.3 kg ­m−3) is characterized by an εNd valueof − 3.3 (Behrens et al. 2018).
The observation of some high εNd values on the OT
side and low values on the Pacific side of the RIA may
result from direct interaction with continental inputs incombination with physical processes, as follows.
(1) Ocean circulation is likely to change near continental zones and shelves owing to local topographical effects
and physical processes. For example, areas with narrow
bathymetry (i.e., the OT) may promote abrupt hydrodynamic behavior that increases horizontal fluxes to values of up to three times vertical fluxes (Huthnance 1995;
Lacan and Jeandel 2005; Lambelet et al. 2016).
(2) Westward-moving eddies to the east of Taiwan
trigger KC meandering. Differing εNd values between
similar locations and depths have been reported in areas

Cruz Salmeron et al. Progress in Earth and Planetary Science(2022) 9:44

Page 11 of 20

Fig. 5 εNd profiles for the study area, together with those reported for the northwestern Pacific (Amakawa et al. 2004b; Wu, 2014; Behrens et al.
2018)

under the influence of mesoscale eddies, which play anessential role in vertical and horizontal mixing processes(Grasse et al. 2012). The KC also undergoes periods ofhigh and low transport. At the southeastern tip of Taiwan, a considerable portion of the KC spreads northeastward during low-transport periods, then flows back to
the OT between the South Ryukyus (Gawarkiewicz et al.
2011). Drifters released to the east of Taiwan indicatecomplex spread patterns within the OT and western PS,
usually reaching the Pacific side of the study area and, toa lesser extent, as far as Okinawa Jima Island (Gawarkiewicz et al. 2011; Rudnick et al. 2015). Particulates discharged by eastern Taiwan rivers likely follow the same
patterns, being deposited on the OT and Pacific sides of
the Ryukyus.

(3) Chemical processes controlling the εNd distribution play a significant role in the variable compositionof the OT and Pacific sides, as discussed in the following
subsections.
5.1.2 Intermediate waterIntermediate water (mainly of 500–700 m depth;σθ = 26.5–27.0 kg ­m−3) on both the OT and Pacific sidesof the RIA has εNd values mainly in the range of − 5.5
to − 4.8, at least 1.7 epsilon units lower than thosereported for NPIW (σθ = 26.67–27.11 kg ­m−3) at similar depths (− 3.1 to − 2.7; Behrens et al. 2018; − 2.8, Wu
2014. Table 2). Although Wu (2014) reported that SCSIWis characterized at similar depths by εNd values ranging from − 4.1 (station LS C3a) to − 3.6 (station SCS C5;
Fig. 1, Table 2), there remains a difference of at least 0.7

Cruz Salmeron et al. Progress in Earth and Planetary Science(2022) 9:44

Page 12 of 20

Fig. 6 εNd values of surface–subsurface (red), intermediate (white), and deep (light blue) waters in the Philippine Sea and the northeastern part of
the South China Sea. Data from Amakawa et al. (2004b), Wu (2014) and Behrens et al. (2018)

epsilon units that cannot be explained by the influence ofSCSIW around the study area.
The presence of a low-εNd source is indicated by lower
εNd values at depths of > 1000 m in the three subareas on the OT side of the RIA (Figs. 3 and 5; e.g., − 8.2
at 1026 m depth in the Yonaguni OT, − 5.1 at 1545 m in
the Iriomote OT, and − 5.9 at 1433 m in the Ishigaki OT).
OT circulation is controlled mainly by Pacific Ocean
water flowing throughout the western margin of the
PS and being carried by the KC from subtropical areas,
with εNd values of surface–intermediate water in the PS
being greater than − 3.7 (Table 2; Behrens et al. 2018).
To a lesser extent, SCS surface–subsurface water also
flows into the OT (Chen 2005), with εNd values of − 4.5
to − 2.9 (Wu 2014; Fig. 6). Even if SCS water fully controlled circulation within the OT, it could be responsible
only for εNd values above − 4.5, which is not the case, so
an external source must provide the lower εNd values.
It is apparent that riverine sediments from Taiwan
contribute to some extent to low εNd values down tothe bottom of the OT. Studies have indicated that up to

60% of the sediment deposited at 1274 m depth at thesouthwestern corner of the OT is derived from Taiwan(Diekmann et al. 2008; Dou et al. 2016), although from an
oceanographic perspective, sediment from Taiwan cannot cross the KC. As the Yonaguni OT is characterized
by the occurrence of intermediate water with the lowest
εNd values of all water columns investigated in this study,some processes (e.g., turbiditic fluxes) must contribute
vertical transport from low-εNd sources to much greater
depths. Sediment fluxes in the southern OT increase
with depth, suggesting horizontal advection in mid-depth
water. Moreover, aluminosilicate plumes identified at
depths of 500–700 m may result from resuspended sediments of the Ilan Shelf and Ridge (Hsu et al. 2004).
5.1.3 Deep waterThe εNd values at 1250–1850 m depth (σθ = 27.5–
27.63 kg ­m−3) in the three subareas on the Pacific sideof the RIA are about − 4 and lie between those of thewestern PS (− 4.1 to − 3.2) and SCS (− 4.4 and − 4.3; stations P2 and SCS C5 of Wu (2014), respectively; Fig. 5).

Cruz Salmeron et al. Progress in Earth and Planetary Science(2022) 9:44

Page 13 of 20

Table 3 δD and δ18O values of water masses on the Okinawa Trough and Pacific sides of the southern Ryukyu Island Arc
Water mass

Depth (m)δ18O (‰)

δD (‰)

Minimum Average Maximum Standard Minimum Average Maximum Standard
deviation
deviation
Okinawa Trough side
Kuroshio Surface Water

< 100

–

4.5*

–

–

–

0.5*

–

–

Kuroshio Tropical Water

100− 200

2.4

3.6

4.8

0.61

0.1

0.4

0.6

0.13

Western North Pacific Central Water

250− 500

1.0

3.1

5.2

1.04

0.1

0.4

0.7

0.14

Kuroshio Intermediate Water

500− 755

− 0.9

1.0

2.9

0.95

− 0.1

0.2

0.4

0.13

Deeper Kuroshio Intermediate Water

755− 1900

Bottom water

1900− 1938 –

− 1.2

0.4

1.9

0.76

–

–

–

− 0.2

–

0.0

0.3

0.11

–

–

–

Pacific side
Surface water

< 83

1.0

3.3

5.7

1.17

0.1

0.4

0.7

0.14

North Pacific Tropical Water

99− 216

1.8

3.4

5.0

0.82

0.2

0.4

0.7

0.13

Western North Pacific Central Water

222− 476

0.7

3.1

5.4

1.17

0.1

0.4

0.7

0.16

North Pacific Intermediate Water

478− 863

1.3

3.5

1.10

0.5

0.17

864− 1851

− 0.7

0.5

1.7

0.62

− 0.2

0.1

Deeper North Pacific Intermediate
Water

− 0.9

− 0.2

0.1

0.3

0.11

Pacific Deep Water

1907− 2620 − 1.0

0.7

2.4

0.86

− 0.1

0.0

0.2

0.08

*Only one data available

Previous studies have shown that εNd values decrease
from the PS toward the SCS, as follows: western PS,
εNd =  − 2.8 at 1500 m (σθ = 27.58 kg ­m−3) at station 5 ofBehrens et al. (2018); Luzon Strait, − 3.5 at 1600 m depth
at station P2; northeastern SCS, − 3.9 at 1600 m depth atstation LS C3a; and − 4.4 at 1600 m depth at station SCS
C5 of Wu (2014; Fig. 5). This confirms that SCSIW flows
out through the Luzon Strait into the PS at 350–1350 m
depth (Wang and Chen 1998). Subsequently, SCSIW
encounters NPIW, resulting in KIW having salinity andεNd values between those of SCSIW and NPIW.
Deep waters in the OT are ventilated by intermediate water from the western PS, which enters via theKG at ~ 1100 m depth (Nakamura et al. 2013). However, due to the limited Nd isotope data available fordepths > 1500 m (σθ > 27.35 kg ­m−3) in the OT, it is difficult to confirm whether its Nd isotope compositionis similar to that of upper deep water (~ 2000 m, σθ= 27.66 kg ­m−3) on the Pacific side of the RIA. Several
factors may explain the low εNd values, as follows.
Our two εNd values for the deepest sites in the OT(− 5.9 at 1433 m, σθ = 27.36 kg ­m−3 in the Ishigaki
OT; − 5.1 at 1545 m, σθ = 27.37 kg ­m−3 in the IriomoteOT) suggest that the upper layer (immediately above
1500 m, σθ < 27.35 kg ­m−3) ventilates and imparts a low
εNd signature to the lower layer (Fig. 3). It is likely thatvertical mixing between intermediate and deeper watershomogenizes the Nd isotope composition below 500 m
depth, as indicated by the uniform εNd values (− 5.9
to − 4.8) in the Iriomote and Ishigaki OT. Moreover,salinity and δD and δ18O values are similar at > 750 m
depth (σθ > 27.0 kg·m−3; Fig. 3). Horibe and Ogura (1968)
reported that considerably high δD values characterize not only surface water but also bottom water aroundthe KC path, compared with adjacent open-sea values(Additional file 5: Fig. S3). They suggested that as the KC
flows northeastward, the influence of surface and subsurface waters with high δD values reaches deeper waters(e.g., intermediate water had δD values of 4.4‰ at 433 m
depth and 3.5‰ at 647 m depth at station 58; Fig. 1). Surface and bottom-water mixing due to a KC-induced flow
may reach > 500 m depth. Our results indicate considerably high δD values in deep waters of the study area, withaverage values of 0.4‰ (σ = 0.8‰) at 1500–1938 m depthon the OT side and 0.7‰ (σ = 0.9‰) at 1907–2620 m
depth on the Pacific side (Fig. 3 and Additional file 5:
Fig. S3). Therefore, the presence of low εNd values evenbelow the intermediate-water domain must be promoted
to some extent by vertical mixing, which decreases withdepth as shown by the differences in δD values below
750 m depth (Fig. 3, Table 3).
Another possible explanation of the negative εNdvalues in the deeper areas is that turbidites, reported
at ~ 1400 m depth to the east of Taiwan in the southern OT (Huh et al. 2006), carry sediment with highly
negative εNd values to deeper areas. In addition, nearcontinents or marginal seas, the εNd distribution ismore complex than that in the open ocean. Therefore,
near the continents, the εNd value of a water masscannot be explained only on the basis of conservativeCruz Salmeron et al. Progress in Earth and Planetary Science(2022) 9:44

behavior and water mass mixing/advection, so local
sources must be involved (Grasse et al. 2012). Boundary exchange and benthic Nd fluxes may affect the εNdvalues of seawater in complex ways (Lacan and Jeandel
2005; Abbott et al. 2015). However, as we do not havedata to determine whether this process is critical, it is
not discussed further here.
Deep-water (2000–3000 m depth, σθ = 27.65–
27.73 kg ­m−3) εNd values decrease from east to west
in the PS from station 5 (− 3.5 to − 2.8) and station 4(− 3.8 to − 3.5) of Behrens et al. (2018) to station P2(− 4.2 to − 3.9) and station WPS N18-05 (− 4.3 to − 3.6)of Wu (2014; Figs. 1, 5). Deep water with low εNd signatures of the western PS then flows northeastward tohigher latitudes along the western PS margin, as shownby data for the Iriomote Pacific (− 5.3; this study) andstation LM-6/11 (− 6.0; Amakawa et al. 2004b). Our
results indicate that the western PS is more affected
by UCDW with low εNd values (i.e., − 6.5 to − 4.6;
Amakawa et al. 2013) than the eastern PS below theintermediate layer, as shown by lower εNd values in
the Iriomote Pacific and Ishigaki Pacific. UCDW flows
in a SE–NW direction, bordering Papua New Guinea.
About 10 × ­106 ­m3 ­s−1 (= 10 Sv) of seawater enters thePS via the East and West Caroline basins and the East
Mariana Basin. Of the total UCDW influx (10 × ­106­m3 ­s−1) into the PS, at least 7.5 × ­106 ­m3 ­s−1 of UCDW
does not flow out of the PS, as indicated by the westernboundary deep-water flowing to the north at 2650 m
depth to the east of Okinawa Jima Island (Kawabe andFujio 2010). Therefore, the remaining flux must be
upwelled, imprinting low εNd signatures on the upper
layers by vertical upwelling and mixing.
5.2 Hydrogen and oxygen isotopes

The seawater δ18O values and CTD salinity measurements provide snapshot data. The CTD-based salinity
ranges are 34.24–34.94 on the OT side and 34.24–34.95
on the Pacific side, whereas the JODC-based salinity (30 yr average) range is 34.29–34.88 on the OT sideand 34.18–34.90 on the Pacific side (Additional file 4:
Fig. S2). Taking into account salinity–seawater δ18O
relationships reported in previous studies (e.g., seawater δ18O = 0.31 × salinity − 10.3 in surface water in theSouth Ryukyus; Abe et al. 2009) and the slight difference
between the CTD- and JODC-based salinity, the measured seawater δ18O values are not much different fromlong-term average seawater δ18O values.
Our δD data are snapshot values rather than long-term
data, such as the JODC salinity and temperature data, sothey may not necessarily represent average δD values ofwater masses. However, as the same is true for most ofthe δD data in previous studies (Benetti et al. 2017; Belem

Page 14 of 20

et al. 2019), it is appropriate to use our δD data here indelineating water masses around the RIA. Therefore, theargument based here on δD data is less rigorous than thatbased on Nd and δ18O data.
On the OT side of the study area, we measured theδD and δ18O values of one sample from KSW in Ishigaki OT. The values were 4.5‰ and 0.5‰ (Additional
file 1: Table S1), with average values for KTW of 3.6‰
and 0.4‰, respectively. On the Pacific side of the RIA,
surface water had average δD and δ18O values of 3.3‰
and 0.4‰, and NPTW had average values of 3.4‰ and0.4‰, respectively (Fig. 3, Table 3). These values indicatethat high δ18O values characterize the KC, as shown inprevious studies (e.g., Horibe and Ogura 1968). Our δDvalues are similar to those reported previously for NWP
waters at low latitudes. For example, Horibe and Ogura(1968) reported δD values of surface water (2.7‰–3.7‰
at < 130 m depth) and tropical water (2.8‰–3.4‰ at 130–
174 m) from east of Luzon Island near the KC domain in
the western PS (station 85, Fig. 1).
The intermediate-water core on the OT side had average δD and δ18O values of 1.0‰ and 0.2‰, which are
lower by ~ 2.1‰ and ~ 0.2‰ than those of WNPCW,
respectively (Table 3). On the Pacific side, the values
were 1.3‰ and 0.1‰, lower by ~ 1.8‰ and ~ 0.3‰ thanWNPCW, respectively. The δD values of the deeper
intermediate water (at > 800 m) on the OT side (0.4‰)
are almost similar to those on the Pacific side (0.5 ‰),although salinity and potential density data (Sect. 2.2.1)
indicate that SCSIW, characterized by lower δD values(e.g., − 0.2‰–0.1‰ at 448–873 m depth) than NPIW in
the western PS (e.g., 0.0‰–0.8‰ at 475–801 m; Horibeand Ogura 1968), enters the OT side. The δ18O values
in the study area are consistent with those reported at
600 m depth near Kerama Jima Island (~ 0‰; Takayanagi
et al. 2012).
The δ18O value of deep water in the NWP (1499 m
depth at 33°N, 146.75°E) is − 0.1‰ (Bigg and Rohling2000), which is within the range of values obtained at
876–1500 m depth on the Pacific side of the study area(− 0.1‰–0.2‰). However, δ18O values are uniformly 0‰on the OT side of the study area at 755–1938 m depth.
On the OT side of the study area, the δD and δ18O values
for KSW (4.5‰ and 0.5‰,) KTW (2.4‰–4.8‰ and 0.1‰–
0.6‰), and WNPCW (1.0‰–5.2‰ and 0.1‰–0.7‰)
overlap, so the three water masses cannot be distinguishedusing these isotopes. This is also the case for the Pacificside of the RIA, where the ranges in δD and δ18O values of
surface water, NPTW, and WNPCW also overlap.
In contrast, average δD and δ18O values differ between
subsurface water (δD ≥ 3.1‰ and δ18O ≥ 0.4‰) andintermediate water (δD ≤ 1.0‰ and δ18O ≤ 0.2‰) on
the OT side of the RIA (Table 3), as is also the case forCruz Salmeron et al. Progress in Earth and Planetary Science(2022) 9:44

Page 15 of 20

Fig. 7 δ18O–δD cross-plots for seawater on the Okinawa Trough (a, b) and Pacific (c) sides of the study area. a Cross-plots for < 433 m depth; b,c plots for seawater collected from all depths. Outliers for which δ18O and/or δD values are not within the mean ± 2σ range for individual water
masses are not shown. For abbreviations, see Fig. 3

subsurface water (δD ≥ 3.1‰ and δ18O ≥ 0.4‰) andintermediate water (δD ≤ 1.3‰ and δ18O ≤ 0.1‰) on the
Pacific side of the RIA (Table 3).
5.3 Relationships among temperature, salinity,and hydrogen and oxygen isotopes

The δD–δ18O cross-plots for surface–subsurface water(Fig. 7) indicate the influence of runoff near continental
areas, with the intercepts representing freshwater inputs(Benetti et al. 2017). The δD and δ18O values of seawater decrease as the influence of freshwater and precipitation increases because freshwater and meteoric water are
depleted in D and 18O. Salinity–δD/δ18O and temperature–δD/δ18O relationships (Additional file 6: Fig. S4)
have been used to characterize water masses, delineate
water-mass mixing, and identify the influence of precipitation/evaporation on surface waters. The three subareas
on the Pacific side of the RIA display similar trends incross-plots for surface to deep waters (Fig. 7c), with noneof the three subareas forming a discrete cluster of values. This reflects a more stable circulation, which could
be attributed to the long-distance high-speed KC and thecomplex fluxes generated. Intermediate water in Ishigaki

Pacific (orange rhombuses) subarea is clustered to the leftof the Iriomote and Yonaguni subareas, as it is characterized by the lowest salinities (Additional file 6: Fig. S4e).
In addition, in Fig. 4, the Ishigaki Pacific corresponds tothe low salinity curve. This correspondence is likely to be
attributed to the difference in salinities between NPIW
and SCSIW. Therefore, the Ishigaki Pacific better retains
the physical properties of NPIW, which is represented
by a distinctive curve with the lowest salinities at about
600–700 m water depth, between σθ = 26 and 27 kg ­m−3.
Intermediate water on the OT side of the RIA is moresaline by ~ 0.11 than that on the Pacific side, for which
there are two possible explanations. First, PDW is moresaline than intermediate water, and as it approaches the
RIA with decreasing depth, it may mix with the intermediate water before entering the OT. However, this is notplausible, because deep water in the OT is topographically isolated from western PDW (UCDW and North
Pacific deep water; Nakamura et al. 2013; Behrens et al.
2018). Although the ventilation process in the OT is notwell understood, σθ at > 1500 m depth on both the OT
and Pacific sides of the RIA (σθ = 27.4 kg ­m−3) indicatesthat the deeper intermediate water flowing from thenorthern PS into the OT via the KG at ~ 1100 m depthCruz Salmeron et al. Progress in Earth and Planetary Science(2022) 9:44

may ventilate as far as the southern OT (at 2000 m depth;
Nakamura et al. 2013). The second and more plausible
explanation for the higher salinity is that the SCSIW(Wang and Chen 1998; Chen 2005), which is modified
KIW becoming more saline in the southern SCS (salinity
34.6; Qu et al. 1999), flows out through the Luzon Strait
at 350–1350 m depth (Wang and Chen 1998) with salinity of 34.4–34.6 (Gong et al. 1992; Qu et al. 2000). The
mixed water comprising SCSIW (55%; wine-colored symbols in Fig. 3) and NPIW (45%) enters the OT via the TY(Nakamura et al. 2013). A salinity front separating western high-salinity water from eastern low-salinity water
has been detected to the east of Taiwan at 21.45°N (Chen
2005), suggesting that SCSIW interacts with NPIW inthis area. The KC has a turbulent flux that causes surface and subsurface waters to reach down to 500 m depth(Horibe and Ogura 1968). Bottom-water upwelling has
been reported from the southern OT (Nakamura et al.
2013). Therefore, the minimum salinity layer within the
OT may increase through vertical mixing with higher
salinity of upper and lower layers.
From surface to bottom water, the Ishigaki OT has
an intercept of 0.82, Iriomote OT 0.27, and Yonaguni
OT 0.15. The very different pattern for Yonaguni OT iscaused by the lack of data for surface water, resulting ina slope considerably lower than those of the Iriomote
and Ishigaki OT (Fig. 7b). The δ18O–δD relationshipat < 433 m depth on the OT side indicates that the Ishigaki OT has the highest intercept (δD = 2.56‰), followed by Iriomote (2.04‰) and Yonaguni OT (0.97‰).
This indicates the more significant influence of riverine
discharge in the Yonaguni and Iriomote subareas thanin the Ishigaki OT (Fig. 7a). Although the proximity of
Yonaguni OT to Taiwan inputs is clear, additional processes must control vertical circulation, as shown by thelow εNd values and similar salinities and δD and δ18O
values below 750 m depth. On the OT side, temperatureand salinity are lower by ~ 1.0 °C–1.5 °C and ~ 0.03–05
in the Yonaguni OT than in the Iriomote OT and Ishigaki OT, respectively. In contrast, the Iriomote OT showshigher salinities at > 250 m depth (Fig. 3). However, Fig. 4
shows similar θ–salinity curves for the three subareas atσθ = 23–25 kg ­m−3. These results suggest that changes inθ and salinity within the three subareas may be related
to vertical variations in physical properties (e.g., a vertical shift of the thermocline). Thermocline shifts are likely
to be driven by mesoscale eddies (Chen and Chen 2020),
which in turn affect KC behavior (e.g., meandering andupwelling) and sea surface height. Areas with sea surface
height anomalies are related to eddies, which promote
geostrophic circulation (Itoh and Yasuda 2009). The presence of eddies from the east coast to the northeast of Taiwan (Zhang et al. 2001; Gawarkiewicz et al. 2011) may

Page 16 of 20

cause vertical shifts in the physical properties of wateraround the Yonaguni OT.
Another possibility is as follows. The cross-plots indicate that the δD and δ18O values of intermediate waterin the Yonaguni OT overlap with those of the upper andlower layers. Such a trend is not confirmed for the Iriomote and Ishigaki OT subareas (Fig. 7c–f ). The interaction of the KC with the topography of the southern OT
may promote cross-shelf and downslope transport,enhancing vertical mixing in this area (Huh et al. 2006).
Therefore, a different circulation pattern should govern
the Yonaguni OT, causing intermediate water to ventilate
down to bottom water.

6 Conclusions
The water-mass structure was delineated on the OT and
Pacific sides of the southern RIA (off Yonaguni Jima, Iriomote Jima, and Ishigaki Jima islands) based on εNd valuesof benthic foraminifers, δD and δ18O values, and physical properties (temperature and salinity) of seawater. δDand δ18O values were determined for seawater collected
at 69–1938 m depth on the OT side and 68–2620 m on
the Pacific side of the study area. εNd values were determined for benthic foraminifers collected at 69–1545 m
depth on the OT side and 74–2264 m on the Pacific sideof the study area. The main results of this study are as
follows.
1. Riverine sediment discharging into the ECS, including the southwestern OT, is derived mainly from old
continental rocks of the Eurasian continent and Taiwan with low εNd values. The low Nd isotope signature is imprinted in water masses flowing within the
OT, resulting in a wide range of εNd values of − 8.2
to − 2.2, with slightly lower values than those of the
Pacific side for seawater at the same depth.
2. The OT-side circulation is disturbed when the highspeed KC interacts with the complex and narrow
bathymetry. As a result, there are overlapping εNdvalues between different seawater layers. The εNdvalues on the OT side suggest that additional processes controlling the Nd behavior of seawater occur
in individual seawater layers. Surface and subsurface waters are influenced by Taiwanese river discharge associated with temporospatial variations inoceanographic conditions such as KC meandering;intermediate water has the lowest εNd values (as low
as − 8.2), possibly resulting from sediment plumes
and turbiditic fluxes; bottom water has εNd values
similar to those of intermediate water, suggesting
active upwelling and vertical mixing.
3. Unlike the OT side, the Pacific side has betterdefined εNd profiles, with more pronounced changes

Cruz Salmeron et al. Progress in Earth and Planetary Science(2022) 9:44

between subsurface–intermediate and intermediate–bottom water boundaries. Surface and subsurface (> 300 m depth) waters are characterized by highεNd values, contrasting with the low values (downto − 7.0) of subsurface–core intermediate water(400–600 m depth). εNd values increase slightlyto − 4.0 below 750 m depth, remaining relatively
constant to about 2000 m depth, before decreasingslightly below that.
4. On the OT side, δD–δ18O relationships suggests that
the Yonaguni, Iriomote, and Ishigaki subareas have different circulation patterns. Furthermore, the surface–subsurface water in the Yonaguni subarea has the lowest δD and δ18O values, attributed to freshwater inputs
from Taiwan.
5. δD and δ18O values are similar at depths of 0–500 m
on the OT and Pacific sides, indicating intense vertical mixing. In contrast, intermediate and bottom/
deep waters are distinguishable from the upper layersby their lower δD and δ18O values.

Abbreviations
CTD profiler: Conductivity–temperature–depth profiler; ECS: East China Sea;
JODC: Japan Oceanographic Data Center; KC: Kuroshio Current; KG: Kerama
Gap; KIW: Kuroshio Intermediate Water; KSW: Kuroshio Surface Water; KTW:
Kuroshio Tropical Water; LREE: Light rare-earth element; NEC: North Equatorial
Current; NPIW: North Pacific Intermediate Water; NPTW: North Pacific Tropical
Water; NWP: Northwest Pacific; OT: Okinawa Trough; TG: Tokara Gap; TS: Tokara
Strait; TY: Strait between Taiwan and Yonaguni Jima Island; PDW: Pacific Deep
Water; PS: Philippine Sea; REE: Rare-earth element; RIA: Ryukyu Island Arc; SCS:South China Sea; SCSIW: South China Sea Intermediate Water; TIMS: Thermal
ionization mass spectrometry; UCDW: Upper Circumpolar Deep Water;
WNPCW: Western North Pacific Central Water.

Supplementary Information
The online version contains supplementary material available at https://​doi.​
org/​10.​1186/​s40645-​022-​00503-5.
Additional file 1: Table S1. Physical properties and isotopic compositionsof seawater on the Okinawa Trough side of the southern Ryukyu Island
Arc.
Additional file 2: Table S2. Physical properties and isotopic compositionsof seawater on the Pacific side of the southern Ryukyu Island Arc.
Additional file 3: Fig. S1. Sites at which temperature and salinity data
during the period from 1 January 1990 to 31 December 2019 wereobtained (Japan Oceanographic Data Center, 2021) from the southern
Ryukyu Island Arc. Solid and open squares represent sites on the Okinawa
Trough and Pacific sides, respectively. Green, blue, and orange symbols
indicate sites in the Yonaguni, Iriomote, and Ishigaki subareas, respectively.
Additional file 4: Fig. S2. Depth profiles of salinity and temperature for
the Okinawa Trough (upper panel) and Pacific (lower panels) sides of thesouthern Ryukyu Island Arc. The profiles of the three subareas are basedon data collected from 1 January 1990 to 31 December 2019 (Japan
Oceanographic Data Center 2021). Gray symbols indicate salinity andtemperature measured during the GK19 cruise using a conductivity–temperature–depth profiler.

Page 17 of 20

Additional file 5: Fig. S3. δD values of deep water below the Kuroshio
Current and in the adjacent open sea. Data from Horibe and Ogura (1968).
Additional file 6: Fig. S4. Salinity–δ18O/δD and temperature–δ18O/δD
cross-plots of seawater on the Okinawa Trough (a–d) and Pacific (e–h)
sides of the southern Ryukyu Island Arc. Outliers for which δ18O and/or
δD values were not within the mean ± 2σ range for the individual water
masses are not shown. For abbreviations, see Fig. 3.
Acknowledgements
We thank the Chief Scientist H. Katayama and onboard scientists of the GK19
cruise for their assistance. We thank Captain H. Yabe and crew of the Kaiyo
Maru No. 1 for their support and assistance aboard their vessel. The manuscript
was significantly improved by the comments and suggestions of A. Oka (editor) and two anonymous reviewers.
Author contributions
ADCS, HT, and YI conceptualized and designed this study. HT, SW, and TI
determined Nd isotope composition. HT and TM undertook H and O isotope
analyses. HUW identified benthic foraminifers. TI conducted sedimentologicalanalysis of seafloor samples. All authors collaborated in the interpretation of
the data and preparation of the manuscript. All authors read and approved
the final manuscript.
Funding
This work was financially supported by the Japan Society for the Promotion of Science KAKENHI (Grant-in-Aid for Scientific Research) Grant number
19H04251 to HT and Frontier Research in Duo (FRiD) of Tohoku University to
Y.I.
Availability of data and materials
Please contact the corresponding author regarding data requests.

Declarations
Competing interests
The authors declare that they have no competing interest.
Author details
1
Institute of Geology and Paleontology, Graduate School of Science, Tohoku
University, Sendai, Japan. 2 Kochi Institute for Core Sample Research, Japan
Agency for Marine-Earth Science and Technology (JAMSTEC), Yokosuka, Japan.
3
Atmosphere and Ocean Research Institute, The University of Tokyo, Kashiwa,
Japan. 4 Geological Survey of Japan, AIST, Marine Geology Research Group/
Institute of Geology and Geoinformation, Tsukuba, Japan.
Received: 20 October 2021 Accepted: 7 August 2022

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