Introduction
Stable isotopes of organic carbon and nitrogen (δ13C, δ15N) and molar carbon to nitrogen (C / N) ratios are natural tracers
frequently used to identify the source and fate of terrestrial organic matter
(OM) in the estuarine and marine environments (Meyers, 1994; Hedges et al.,
1997; Goñi et al., 2014; Selvaraj et al., 2015). This approach is based
on the significant difference in δ13C, δ15N and
C / N ratios between different end-members (e.g. terrestrial and
marine), and the assumption that only a physical mixing of OM from
compositionally distinct end-members occurs in these marginal settings
(Thornton and McManus, 1994; Hedges et al., 1986). Quantifying the relative
contributions of end-members using mass balance models thus requires known
and constant elemental and isotopic values of end-members and major sources
of OM in the study region (e.g. Goñi et al., 2003). Therefore,
application of mixing models for the discrimination of OM sources requires
clearly identified representative values for local OM sources. However, in
most cases, end-member values of δ13C, δ15N and molar
C / N ratios are represented by “typical” numbers, such as ca. -20
and -27 ‰ for δ13C of marine phytoplankton and
terrestrial plants (Kandasamy and Nagender Nath, 2016, and references therein),
respectively, but without measuring discrete end-member values in real, local
or regional OM source materials. For example, a number of earlier studies
failed to measure isotopic values of marine phytoplankton despite using
end-member mixing models to distinguish marine versus terrestrial OM in
surface sediments (e.g. Kao et al., 2003; Wu et al., 2013), or these numbers
were simply represented by values of particulate organic matter (POM) in surface
waters in the studied system (e.g. Zhang et al., 2007) or elsewhere from
other ocean basins (e.g. Hale et al., 2012). It is known that stable
isotopes (δ13C, δ15N) and molar C / N ratios of POM
in estuarine and marine areas are representative of primary
production-derived OM when POM are mostly derived from phytoplankton biomass
(Gearing et al., 1984). Since phytoplankton are the main primary producer of
marine OM, the elemental and isotopic compositions of phytoplankton should
therefore be considered while studying the dynamics of POM in the marine
water column.
Chlorophyll a (Chl a) concentration in seawater is often used as an
index of phytoplankton biomass (Cullen et al., 1982).
The deep chlorophyll maximum (DCM) layer, which contributes significantly to
the total biomass and primary production in the whole water column (Weston et
al., 2005; Sullivan et al., 2010), is approximately
equal to the subsurface biomass maximum layer (e.g. Sharples et al., 2001;
Ryabov et al., 2010). The formation of maximum chlorophyll concentration at
the DCM layer has been explained by several mechanisms: the differential
zooplankton grazing with depths (Riley et al., 1949; Lorenzen, 1967; Pilati
and Wurtsbaugh, 2003), adaption of phytoplankton to light intensities or to
increased concentration of nutrients (Nielsen and Hansen, 1959; Gieskes et
al., 1978; Hickman et al., 2012), chlorophyll accumulation by sinking
detritus of phytoplankton (Gieskes et al., 1978; Karlson et al., 1996),
decomposition of chlorophyll by light (Nielsen and Hansen, 1959), and
wind-driven nitrate supply and nitrate uptake in seasonally stratified shelf
seas (Hickman et al., 2012; Williams et al., 2013). The DCM layer is common
in both coastal and open oceans, occurring at relatively shallow depths
(1–50 m) in coastal seas, but at greater depths (80–130 m) in the open ocean
(Cullen et al., 1982; Gong et al., 2015), and often variable in time and space
(Karlson et al., 1996). For example, the DCM layers were reported at depths
of 30–50 m across the shelf in the southern East China Sea (ECS) during summer
from 1991 to 1995 (Gong et al., 2010). Hence, δ13C, δ15N
and molar C / N ratios of POM in the DCM layers of the continental shelf
waters should reflect the δ13C, δ15N and molar C / N
ratios of phytoplankton (Savoye et al., 2003, 2012; Gao et al., 2014).
The East China Sea is one of the largest marginal seas in the world, receiving
huge quantities of freshwater (905.1 km3 yr-1; Dai et al., 2010)
and organic C (2.93 Tg C yr-1, Tg = 1012 g; Qi et al.,
2014) from the Yangtze River (Changjiang). Nutrient-rich freshwater inputs in
turn stimulate the water column productivity in coastal waters compared to
the open ocean. Annual primary production over the entire shelf of the East
China Sea is high relative to other marginal seas and was estimated to be
85 Tg C yr-1 in 2008 (Tan et al., 2011). Several studies have been
carried out on the physical, chemical and biological aspects of the East
China Sea, including distributions of seasonal currents (e.g. Gong et al.,
2010), chemical hydrography and nutrient distribution (Chen, 1996, 2008), and
phytoplankton species composition in the water column (e.g. Zheng et al.,
2015; Jiang et al., 2015). Likewise, δ13C, δ15N and
molar C / N ratios of POM have been determined for a limited number of
transects across the East China Sea (e.g. Wu et al., 2003, 2007a) as well as
in a wide area of the western North Pacific marginal seas (Chen et al.,
1996). Nonetheless, studies on elemental ratios and stable isotopic
compositions of POM in DCM layers in the continental shelf of the East China
Sea, especially along the indirect transport pathway of the Yangtze-derived
terrestrial material to the Okinawa Trough (Chen et al., 2017), are poorly
studied. A recent study in the northern East China Sea investigated elemental
and isotopic compositions of POM in the surface, DCM and bottom layers on
both seasonal and inter-annual timescales (Gao et al., 2014); however, there
was minimal attention given to biogeochemical processes associated with the
DCM. Here, we investigate δ13C, δ15N and molar C / N
ratios of suspended POM around the DCM layer in the continental margin of the
East China Sea, in particular the area south of the Yangtze estuary, aiming
(1) to comprehend the sources of POM in DCM layers and (2) to understand the
factors controlling δ13C and δ15N dynamics in DCM layers
of the southern East China Sea.
Study area
The East China Sea (Fig. 1) is the largest river-dominated marginal sea
in the north-western Pacific region (Chen et al., 2017). The ECS shelf is
wide (> 500 km), but relatively shallow (< 130 m) with an average
water depth of 60 m (Gong et al., 2003; Liu et al., 2006). The Yangtze River
(Fig. 1), with a catchment area of more than
1.94 × 106 km2 (Liu et al., 2007), is the main source of
freshwater and sediment to the continental shelf. It is the fifth largest
river in terms of water discharge (900 km3 yr-1) and the
fourth largest river in terms of sediment discharge (470 Mt yr-1)
in the world (Milliman and Farnsworth, 2011).
Map showing the locations of suspended particulate
matter (SPM) collected around the deep chlorophyll maximum (DCM) layer from
the East China Sea during summer (22 June–21 July) 2013 for the present
investigation. Also shown is the modern current pattern in the East China
Sea. Red circles mark the SPM samples that were collected from the water
depths either below or above but mostly contiguous to the DCM layer. CDW –
Changjiang Diluted Water, CCC – China Coast Current, TWC – Taiwan Warm
Current and KC – Kuroshio Current. The dashed ellipse represents the center
of Kuroshio upwelling, occurring due to an abrupt change in the bottom
topography, in the north-east of Taiwan Island (Wong et al., 2000). Also
shown is the PN transect, a cross-shelf transect in the East China Sea in which particulate organic matter dynamics is relatively well studied.
In addition to the huge inputs of nutrients (dissolved inorganic
nitrogen, DIN: 61.0 ± 13.5 × 109 mol yr-1 for
the interval of 1981–2006; Chai et al., 2009) and sediments from the Yangtze
River, the ECS is characterized by a complex circulation pattern that is
largely driven by the seasonally reversing East Asian monsoon winds (He et
al., 2014; Chen et al., 2017). The surface circulation in the shelf is
characterized by the south–north China Coastal Current (CCC) in the west,
northward-moving Taiwan Warm Current (TWC) in the central part and the
north-eastward-flowing Kuroshio Current (KC) in the east (Fig. 1) (Liu et
al., 2006). The Changjiang Diluted Water (CDW) is a mixture of Yangtze River
freshwater and the East China Sea shelf water and is characterized by a low
salinity (< 30, Umezawa et al., 2014). Owing to a huge amount of
freshwater discharge from the Yangtze into the ECS, it is thought that the
CDW is the main component of CCC (Fig. 1). Because of the East Asian monsoon,
where there is a strong north-east monsoon in winter and a weaker south-west
monsoon in summer, the CDW flows southward along the coastline of mainland
China as a narrow jet in winter (Chen, 2008; Han et al., 2013), whereas the
same spreads mainly to the north-east in summer (Isobe et al., 2004). The
TWC is a mixture of the warm water from the Taiwan
Strait and intruding saline Kuroshio water; the latter is thought to be the
most dominant source of heat and salt to the ECS (Su and Pan, 1987; Zhou et
al., 2015). In addition, Kuroshio Subsurface Water (KSSW) is upwelled in the
north-east off Taiwan Island due to an abrupt change in seafloor topography at
the ECS outer shelf (dashed ellipse in Fig. 1) (Su et al., 1990; Sheu et al.,
1999). The upwelled, oxygen-undersaturated KSSW is characterized by low
temperature, high salinity and high nutrients (Liu et al., 1988; Wong et al.,
1991). The water exchange rate between the ECS water and Kuroshio water was
estimated to be about 22 000 ± 9000 km-3 yr-1, which is
approximately 25 times the amount of Yangtze run-off into the ECS (Li,
1994; Sheu et al., 1999). Furthermore, Kuroshio water accounts for up to
90 % of the shelf water in the ECS (Chen, 1996; Sheu et al., 1999).
Summary statistics of elemental and isotopic compositions, as well
as C / N and POC / Chl a ratios, of suspended particulate matter
(SPM) around DCM layers in the southern East China Sea (n=36). Chl a is
the converted value using the linear relationship between measured Chl a
and chlorophyll fluorescence. SD = standard deviation.
Sampling depth
SPM
POC
PN
δ13CPOC
δ15NPN
C / N
POC / Chl a
(m)
(mg L-1)
(µg L-1)
(µg L-1)
(‰)
(‰)
molar
(g g-1)
Minimum
10
1.7
20.4
4.4
-25.8
3.8
4.1
33.3
Maximum
130
14.7
263.0
52.8
-18.2
8.0
6.3
303.3
Mean
45
4.4
85.5
17.7
-23.0
6.1
5.6
100.3
SD
21
2.7
49.5
9.9
1.5
1.0
0.5
51.8
The primary productivity in the ECS is limited by nitrogen in summer, but
light in winter (Chen et al., 2001; Chen and Chen, 2003). With the highest
primary production during summer, annual primary production showed distinct
spatial and temporal variations of 155, 144 and
145 g C m-2 yr-1 in the north-western ECS, south-eastern ECS
and the entire ECS, respectively, in 1998 (Gong et al., 2003). The primary
productivity has, however, decreased by 86 % between 1998 and 2003 due to a
large number of impoundments in the drainage basin of Yangtze River (Gong et
al., 2006).
Materials and methods
Sample collection
To investigate the biogeochemical characteristics of POM in the DCM layer of
the southern East China Sea, suspended particles around the DCM water depths
(10–130 m; Table 1) were collected from 36 stations along 7
transects across the continental shelf by the Science 3 cruise
during summer (22 June–21 July) 2013 (Fig. 1). At each site, the physical
properties of the water column were recorded by a
conductivity–temperature–depth (CTD) rosette (Seabird, SBE911+) fitted with
a Seapoint chlorophyll fluorometer to detect the fluorescence maximum (see
Table S1 in the Supplement for the whole dataset). Seawater was collected
using the rosette of Niskin water bottles attached to the CTD frame, and
then stored in 5 L PVC bottles. All PVC bottles had been soaked in 0.1 M
HCl and then cleaned by distilled water. The volume of each water sample was
measured by graduated cylinder before filtration. Suspended particles were
obtained by filtering 4.1–19.1 L of seawater collected around the
fluorescence maximum layer through 0.7 µm/47 mm Whatman glass
fiber filters (GF/F), which were wrapped in aluminium foil. The filtration
was under an ultimate pressure of 0.08 MPa to avoid the rupturing of
phytoplankton cells (Steinman et al., 2017). All filters had been
pre-combusted at 450 ∘C for 4 h in a muffle furnace to remove the
background carbon and pre-weighed to determine the concentration of
suspended particulate matter (SPM). After filtration, filters were folded
without rinsing, wrapped again in aluminium foil and then immediately stored at
-20 ∘C in a freezer on-board before they were brought
back to the laboratory for further analysis.
Determination of SPM concentration and analyses of Chl
a, POC, PN, δ13C and δ15N
In the laboratory, filters with suspended particles were freeze-dried and
then dried in an oven at 50 ∘C for 48 h. The weight difference
between the dried filter and the same filter before the filtration was used
to calculate the weight of SPM. Five SPM samples (DH1-2, DH2-1, DH3-1, DH7-1
and DH7-7; Fig. S1) from water depths ranging between 20 and 50 m were
randomly selected for the measurement of chlorophyll a (Chl a)
concentration. Chlorophyll a was extracted using 90 % acetone and then
determined spectrophotometrically according to Lorenzen (1967) and Aminot and
Rey (2000). Briefly, the absorbance of sample extraction was measured at 665
and 750 nm against a 90 % acetone blank before (E665o,
E750o) and after (E665a, E750a) acidification
with 1 % HCl by the UV-Vis spectrophotometer (UV 1800, Shimadzu). Chl a
concentration (µg L-1) was calculated as follows: Chl a=11.4 × 2.43 × ((E665o - E750o) - (E665a - E750a)) × Ve/L ×Vf, where
Ve and Vf were the volumes of sample extraction and filtered seawater (mL), respectively, and L was the cuvette light path (cm)
(Aminot and Rey, 2000).
Prior to the measurement of POC and PN contents and their stable isotope
values (δ13CPOC and δ15NPN) in SPM
samples, a half of each filter was placed in a culture dish and 3 mL of 1 N
HCl was then added into the dish by a dropper. These were allowed to react for
16 h to remove inorganic carbon (mainly carbonate). The de-carbonated sample was
dried at 50 ∘C for 48 h in an oven for HCl evaporation. Then a half
of the de-carbonated filter (i.e. a quarter of the original filter,
∼ 11 mm) was then punched and placed in tin capsules for further
analysis. The POC and PN contents and their δ13CPOC and
δ15NPN compositions were measured at the Stable Isotope
Facility of University of California Davis in the USA, by using an elemental
analyser (EA) (Elementar Analysensysteme GmbH, Hanau, Germany) interfaced to
a continuous flow isotope ratio mass spectrometer (IRMS; PDZ Europa 20–20,
Sercon Ltd., Cheshire, UK). During the isotope (δ13CPOC
and δ15NPN) analyses, different working standards (bovine
liver, glutamic acid, enriched alanine and nylon 6) that were compositionally
similar to the samples were used and were calibrated against NIST Standard
Reference Materials (IAEA–N1, IAEA–N2, IAEA–N3, USGS–40 and USGS–41).
The standard deviation was 0.2 ‰ for δ13C and
0.3 ‰ for δ15N. Isotopic values were presented in standard
δ-notation as per mille deviations relative to the conventional
standards, i.e. VPDB (Vienna Pee Dee Belemnite) for carbon and atmospheric
N2 for nitrogen, that is δX
(‰) = [(Rsample-Rstandard)/Rstandard]×103, where X=13C
or 15N, R=13C/12C or 15N/14N, Rsample
and Rstandard are the heavy (13C or 15N) to light
(12C or 14N) isotope ratios of sample and standard, respectively
(e.g. Selvaraj et al., 2015).
Vertical distributions of temperature and salinity along
seven transects across the southern East China Sea during summer 2013. Note
that there is an obvious thermally stratified water column during the
collection of suspended particles in the study area.
Lorrain et al. (2003) cautioned that the measurement of PN and δ15N after freezing increases the uncertainty of δ15N and, in
combination with the concentrated HCl treatment, leads to a loss of PN and
alteration of the δ15N signature. Therefore, PN content and
δ15N values in the current study may have some bias due to
de-carbonation. Nonetheless, a similar methodological approach has been adopted
by Wu et al. (2003) while investigating suspended particles along the PN
transect in the East China Sea (Fig. 1) and by Hung et al. (2007) while
studying the suspended particles in the entire East China Sea. For instance,
the range of δ15N values (∼ 3.8–8.4 ‰) obtained in
the present study is comparable to the range of δ15N values (ca.
0.7–9.4 ‰) obtained by Wu et al. (2003) for the entire water
column. In addition, precision for δ13C and δ15N
decreases for samples containing less than 100 µgC and
20 µgN, respectively. Among 36 filters analysed for the
present study, only 5 (3) filters contain less than 100 µgC
(20 µgN).
Discussion
Influence of different water masses in the southern ECS
In order to identify the different water sources in the study area,
temperature–salinity (T–S) diagrams were drawn for the entire water column
(Fig. 6a) as well as for the SPM sampling depth around DCM layers (Fig. 6b).
The T–S diagram for all the water depths shows a convergence at around
17 ∘C, 34.6 (Fig. 6a), representing the upwelling of KSSW (Umezawa
et al., 2014). There are two trends in the T–S diagram, indicating a mixing
of three water masses: one is less saline and much colder water, mainly CDW,
another is more saline and warmer, mainly Taiwan Warm Current Water (TWCW),
and the third one is KSSW (Fig. 6a). The low salinity observed at five
coastal sites (DH1-1, DH2-1, DH2-2, DH3-1 and CON02; Fig. 2) indicates the
influence of CDW mostly in surface water, but also some of the DCM depths
where water was sampled for SPM. This is also evident from Fig. 6b where five
stations fall within the area of shelf mixed water (SMW), which is a water
body composed of a mixing between CDW and KSSW. However, apart from at these five
coastal stations, most DCM depths where water was sampled for SPM seem to be
weakly influenced by the CDW (Fig. 6b). Based on the T–S range of
different water masses (Fig. 6), we further delineated the area and water
depths influenced by three important water masses: CDW, TWCW and KSSW
(Fig. 7). Interestingly, the influence of CDW was constrained to the upper
10 m in five coastal stations, whereas TWCW influenced the upper 30 m and
covered three quarters of the study region, with KSSW largely influencing the
bottom water across the entire study region (Figs. 2, 6a and 7).
A diagram delineating the regions influenced by three
main water masses based on the T–S relationship (Figs. 2 and 6) in the study
area. Area with grey polygon represents the influence of CDW, which is
limited only in the upper 10 m. The sky-blue area represents the dominance
of TWCW, which is limited to ∼ 30 m below the surface. The
polygon coloured by deep blue represents the area influenced by the KSSW,
indicating that the bottom water of the entire study area was dominated by
KSSW.
Bi-plots showing the relationships of (a) POC vs. PN and
(b) POC vs. Chl a in suspended particulate matter investigated in this
study. Redfield ratio (dashed line in panel a) is taken from Redfield (1958).
In summary, although the river run-off was huge, the influence of CDW plume
in the southern part of the ECS was weak during summer 2013 mainly because
most of the CDW plume was transported north-eastward of the Yangtze estuary
to the Korean coast (Isobe et al., 2004; Bai et al., 2014; Gao et al.,
2014). This contrasts with summer 2003 when the plume front moved southward
(Bai et al., 2014). Meanwhile, the intrusion of TWCW and KSSW was strong in
the continental shelf of the East China Sea during summer 2013.
Characterization of POM in DCM layers
Molar C / N ratio
A necessary first step in the source analysis of POM using bulk carbon and
nitrogen isotopes as well as the molar carbon to nitrogen ratio is to
identify the form of total nitrogen in the measured SPM, so that inorganic
nitrogen is not mis-assigned as nitrogenous organic end-member (Hedges et
al., 1986). The linear relationship between POC and PN (R2=0.98,
p < 0.0001; Fig. 8a) suggests that nitrogen is strongly associated with
organic carbon. The slope of linear regression of POC against PN corresponds
to a molar C / N ratio of 5.76 (Fig. 8a). The positive intercept on the
PN axis when POC is zero represents the amount of inorganic nitrogen
(∼ 0.03 µM), indicating that essentially all nitrogen is in
organic form. The molar C / N ratios of all SPM samples (4.1–6.3)
from the DCM layers are lower than the canonical Redfield ratio (6.63)
(Fig. 8a), but are similar to the average molar C / N ratios of 5.6 for
marine POM (Copin-Montegut and Copin-Montegut, 1983) and 6 for POM in cold,
nutrient-rich waters at high latitudes (Martiny et al., 2013). The range also
falls within the range of 3.8 to 17 reported for marine POM (Geider and La
Roche, 2002), but it is higher than an unprecedented low C / N ratio
(2.65 ± 0.19) of POM in the Canada Basin that was attributed to a dominant
contribution of smaller-sized (< 8 µm) phytoplankton to POC
(Crawford et al., 2015). Wu et al. (2003) investigated the C / N ratio of
POM (4.3–29.2) at all depths along the PN transect, a standard
cross-shelf section extending from the Yangtze estuary south-east to the
Ryukyu Islands, cross-cutting the Okinawa Trough and perpendicular to the
principle axis of Kuroshio Current in the ECS (Fig. 1). Liu et al. (1998)
measured the C / N ratio of POM in the surface water of the ECS and found
a wider C / N ratio from 4.0 to 26.9 with a mean ratio of 7.6 in spring
and from 4.7 to 34.3 with a mean ratio of 15.2 in autumn 1994. The authors
attributed the lower C / N in spring to more intense biological activity
than in autumn, and the spatial distribution of C / N was thought to be
related to that of phytoplankton abundance.
Characteristically, a narrow range of low C / N ratios in our SPM samples
and less influence of CDW in the study region (Fig. 7) confirm the lack of
terrestrial signals transported mainly by the Yangtze River. We therefore
suggest that the POM in the DCM layers of the southern East China Sea is
dominated by marine-sourced OM with an unrecognized contribution of
terrestrial OM. Low C / N ratios further restrict the assumption of
degradation of nitrogen-rich OM, a process that normally increases the
C / N ratio to more than that of the Redfield ratio. Therefore, the molar
C / N ratio can be better explained as a source signal of OM rather than
OM degradation in the SPM investigated in this study.
POC / Chl a ratio
The linear correlation between POC and Chl a (R2=0.49, p<0.0001;
Fig. 8b) further indicates that the phytoplankton productivity is largely
responsible for the POC production in the SPM samples. Moreover, the
POC / Chl a ratio of 34.1 g g-1 derived from the slope of a
regression line (y=34.1(±9.99)x+49.9(±8.86)) (Fig. 8b) is consistent
with the reported POC / Chl a ratios in the ECS (36.1 g g-1;
Chang et al., 2003) and the north-western Pacific (48 g g-1; Furuya,
1990). However, the POC / Chl a ratio obtained in this study is lower
than that estimated (64 g g-1) for the sinking particles in the ECS
and the Kuroshio region, off north-eastern Taiwan Island (Hung et al., 2013). The
range is well within the range (13–93 g g-1) reported for POM in the
ECS by Chang et al. (2003) and is also consistent with the range
(18–94 g g-1) estimated from phytoplankton cell volumes by the same
authors. Although the Chl a concentration in our study was converted based
on the linear relationship between measured Chl a and in situ fluorescence
values (see Sect. 3.2 and Fig. S1 in the Supplement for more details), it is
more or less similar to Chl a concentrations obtained in the
above-mentioned studies, which were mostly extracted from filtered particles
(Chang et al., 2003; Hung et al., 2013).
The POC / Chl a ratio has been used for the discrimination of POM sources
in coastal ocean waters (Cifuentes et al., 1988). The POC / Chl a ratio in
living phytoplankton varies with temperature, growth rate, day length,
phytoplankton species and irradiance (Savoye et al., 2003 and references
therein). The POC / Chl a ratio of living phytoplankton was reported to
be between 40 and 140 g g-1 (Geider, 1987; Thompson et al., 1992;
Montagnes et al., 1994; Head et al., 1996). Furthermore, a POC / Chl a
ratio of less than 200 g g-1 is an indication of a predominance of
newly produced phytoplankton (or autotrophic-dominated) in POM, and that a
value higher than 200 g g-1 is an indication of detrital or degraded
organic matter (or heterotrophic/mixture-dominated) (Cifuentes et al., 1988;
Savoye et al., 2003; Liénart et al., 2016). The POC / Chl a ratio
in the DCM layer of the ECS is almost < 200 g g-1
(33–200 g g-1), with one exception (CON02: 303 g g-1; Fig. 9),
indicating that POM in the DCM layers of ECS was dominated by phytoplankton,
as also indicated by the low C / N ratios (4.1–6.3). The relatively high
POC / Chl a ratio in only one station, CON02 (Fig. 9), suggests that the
POM in this sample was likely sourced from degraded phytoplankton OM,
terrestrial OM or heterotrophic-dominated OM. However, the molar C / N
ratio of CON02 (5.3) is lower than the canonical Redfield ratio (6.63),
eliminating the probability of degraded and terrestrial OM sources. In
addition, the insignificant linear correlation between the C / N ratio and
POC / Chl a ratio (Fig. 9) supports the non-degraded POM, a process
resulting in a simultaneous increase in C / N and POC / Chl a
ratios, mainly because of the preferential decomposition of N-rich OM, as
well as a fast degradation of Chl a than the bulk POC pool (e.g. Savoye et
al., 2003). Thus, the POM in CON02 seems to be dominated by heterotrophic
biota, though the exact reason for the dominance of heterotrophic biota
at only one location in our study area is unknown and needs further investigation.
Molar C / N ratio vs. POC / Chl a ratio in suspended
particulates investigated in this study. The vertical line represents
POC / Chl a ratio of 200 g g-1, the upper limit for
phytoplankton-dominated particulate organic matter (Savoye et al., 2003). See
text for more details. CON02 is the station where red tide was observed
during the sampling time and the colour of the surface water was brown and
dissolved oxygen in the bottom water was 1.6 mg L-1.
Briefly, several clues indicate the predominance of newly produced,
phytoplankton-synthesized OM around DCM layers of the southern East China
Sea: (1) low influence of fresh water, (2) low molar C / N ratios, (3) a
linear correlation between POC and chlorophyll a, and (4) low POC / Chl
a ratios, mostly < 200 g g-1.
Bi-plots showing the relationships of (a) δ13CPOC vs. POC and (b) δ15NPN
vs. PN in suspended particulate matter around the deep chlorophyll maximum
layer in the southern East China Sea.
Dynamics of δ13CPOC in POM in DCM
Although a narrow range of molar C / N ratio in the SPM indicated an
aquatic origin for the POM at DCM layers, the wide variability of δ13CPOC (-25.8 to -18.2 ‰) suggests that the POM
around DCM layers would be a mixture of terrestrial C3 plants with a typical
δ13C value of ca. -27 ‰ (e.g. Peters et al., 1978; Wada
et al., 1987) and marine phytoplankton with a typical δ13C range of
-18 to -20 ‰ (e.g. Goericke and Fry, 1994). However, Fig. 5
illustrates a distinct decreasing trend in δ13CPOC
towards the outer shelf; a pattern opposite to an increasing trend in
δ13C evident in suspended particles and surface sediments, i.e.
seaward decrease in terrestrial OC in surface sediments of many
river-dominated margins (Emerson and Hedges, 1988; Meyers, 1994; Hedges et
al., 1997; Kao et al., 2003; Wu et al., 2003). Such a spatial distribution
with less negative δ13CPOC values in the coastal
region, but more negative δ13CPOC values in the
middle-outer shelf is inconsistent with the idea of terrestrial OC influence.
The elevated δ13CPOC values (average of
-20.7 ‰) in the coastal region, concomitant with high POC
concentrations (Fig. 4), are consistent with the higher marine primary
productivity (11 g C m-2 yr-1) reported in the western than
that in the eastern parts of the East China Sea (Gong et al., 2003). The lower
δ13CPOC occurred in the middle–outer shelf region where
oligotrophic Taiwan Warm Current Water and Kuroshio Water spread (Fig. 5).
The lowest δ13CPOC (-25.8 ‰) was observed at
a water depth of 85 m, off north-eastern Taiwan, likely due to the intrusion of
Kuroshio Subsurface Water with low δ13C from -31 to
-27 ‰ Wu et al., 2003), is also in agreement with the hydrographic
parameters of this location (Figs. 2 and 7).
A positive linear correlation between δ13CPOC and POC
(R2=0.55, p<0.0001; Fig. 10a), a characteristic feature of
productive oceanic regions (Savoye et al., 2003), suggesting the effect of
growing primary productivity (and or increasing cell growth rate) on a
decrease in carbon fractionation during photosynthesis (Miller et al., 2013).
This is likely because of a limitation of dissolved CO2, which cannot be
compensated for in time by the surrounding water in a relatively closed system
because of stratification (Kopczyńska et al., 1995). Further, high
productivity makes 13C-enriched OM in phytoplankton (Fry and Wainwright,
1991; Nakatsuka et al., 1992; Miller et al., 2013). Lowe et al. (2014)
observed increased δ13C and fatty acid concentration in the POM
while increasing phytoplankton abundance in the nearshore waters of San Juan
Archipelago, WA. Although primary productivity has a significant correlation
with δ13CPOC, only 55 % of δ13CPOC variation can be explained by primary productivity
(Fig. 10a), implying that other factors, such as species and sizes of
phytoplankton, must have influenced δ13C values of phytoplankton
living in the DCM layers (Falkowski, 1991; Hinga et al., 1994).
The distribution of the phytoplankton community in the East China Sea is affected
by physico-chemical properties (temperature, salinity and nutrients) of
different water masses and surface currents (Umezawa et al., 2014; Jiang et
al., 2015, 2017). Diatoms and dinoflagellates are the main phytoplankton
communities in summer, with 136 taxa of diatoms from 55 genera and 67 taxa of
dinoflagellates from 11 genera having been reported, along with minor
communities of chrysophyta, chlorophyta and cyanophyta (Guo et al., 2014b).
There is a clear decreasing trend in phytoplankton abundances in the East
China Sea from the surface to the bottom, as well as from the coastal to offshore
region that is widely believed to be due to nutrient availability (Zheng et
al., 2015). The phytoplankton species have distinct spatial characteristics,
but no significant differences in species between surface waters and the DCM
layers (Zheng et al., 2015). Diatoms with large cell sizes were the dominant
species in the coastal region, while phytoplankton with small sizes was
dominant in the oligotrophic offshore shelf and Kuroshio waters (Furuya et
al., 2003; Zhou et al., 2012). According to Jiang et al. (2015), the
contribution of micro- (> 20 µm), nano- (3–20 µm) and
pico-phytoplankton (< 3 µm) to Chl a was 40, 46
and 14 %, respectively, in nutrient-rich inshore waters, and 14, 34 and 52 % in
offshore regions in summer 2009. The outer shelf region was composed of small-sized phytoplankton, mainly cyanobacteria and cryptophytes transported by
Taiwan Warm Current and Kuroshio Current. It has been reported that diatoms
have higher δ13C values (-19 to -15 ‰) than
dinoflagellates (-22 to -20 ‰; Fry and Wainwright, 1991; Lowe et
al., 2014). Likewise, large phytoplankton have higher δ13C values
than small phytoplankton and heterotrophic dinoflagellates have higher
δ13C values than autotrophic dinoflagellates (Kopczyńska et
al., 1995). Similarly, wide variations of δ13CPOC
(-22.05 to -27.62 ‰) at DCM layers in the northern East China
Sea were documented by Gao et al. (2014). Significant variations of δ13C in suspended OM that was dominated by phytoplankton were reported
from the Delaware estuary (-25 to -20 ‰; Cifuentes et al.,
1988), the Bay of Seine (-24.3 to -19.7 ‰; Savoye et al., 2003),
the Santa Barbara Channel (Miller et al., 2013) and the nearshore waters of
San Juan Archipelago, WA (-24.1 to -18.9 ‰; Lowe et al., 2014).
These variations were influenced largely by the isotopic fractionation during
phytoplankton photosynthesis and degradation than by changes in the relative
contributions of terrestrial and aquatic OM (Fogel and Cifuentes, 1993;
Savoye et al., 2003).
Bi-plots showing the relationships of (a) δ13CPOC vs. temperature for samples separated into two groups
based on temperature: < 24 ∘C and > 24 ∘C,
(b) temperature-normalized δ13C (δ13C24∘C) vs. POC concentration,
(c) δ13C24∘C vs. POC / Chl a
ratio, and (d) δ13C24∘C vs. molar
C / N ratio in suspended particulate matter around deep chlorophyll
maximum layers in the southern East China Sea.
Temperature effect on the δ13CPOC around the DCM layer
Apart from primary production and the growth rate and species composition,
temperature and biomass degradation may influence the carbon isotopic
composition of phytoplankton (Savoye et al., 2003). Temperature has an
indirect effect on isotopic fractionation between phytoplankton carbon and
dissolved CO2, and therefore on phytoplankton δ13C (e.g. Rau
et al., 1992; Savoye et al., 2003). The C / N ratio, POC / Chl a
ratio and δ13CPOC all indicated that the POM around the
DCM layer is dominated by newly produced phytoplankton OM (see
Sect. 5.1–5.3). Therefore, to understand the temperature effect on
δ13C of phytoplankton, we plotted our δ13CPOC
data against temperature after separating samples into two groups based on temperature: < 24 ∘C and > 24 ∘C (Fig. 11a). Data points of both groups show a
decreasing δ13C of phytoplankton biomass with increasing
temperature around the water depths of DCM in the southern ECS (Fig. 11a).
Such a relationship is in contrast to the positive relationship between these
two variables observed for the surface ocean POM around the world (Sackett et
al., 1965; Fontugne, 1983; Fontugne and Duplessy, 1981).
The negative relationship between δ13CPOC and
temperature is likely related to biological activity and carbonate
dissolution equilibrium; both may control the concentration of dissolved
inorganic carbon in the DCM layers, which are closer to euphotic depths (see
Sect. 4.1). The weak correlation between δ13CPOC and
temperature supports a weak influence of temperature on δ13CPOC around DCM layers in the study area (Fig. 11a). A
decrease in fractionation of approximately
-0.56 ‰ ∘C-1 is estimated for POM collected at
< 24 ∘C, whereas a decrease in fractionation of roughly
-0.51 ‰ ∘C-1 is estimated for POM collected at
> 24 ∘C (Fig. 11a). In order to distinguish the influence of
biological parameters from temperature on δ13CPOC, the
δ13CPOC data were corrected for the “temperature
effect” by normalizing the data using the following equation:
δ13CPOC=f(T).
In the present study, since most δ13CPOC values come
from the DCM layer and the δ13CPOC is negatively
correlated with temperature (Fig. 11a), we applied our own temperature
coefficients (-0.56 and
-0.51 ‰ ∘C-1) and δ13CPOC was
normalized at 24 ∘C (i.e. the mean temperature at sampled water
depths) using the following formula (Savoye et al., 2003):
δ13C24∘C=δ13CPOC-s(T-24), where δ13C24∘C is the
temperature-normalized δ13CPOC, T is the seawater
temperature in ∘C from water depths where SPM sampled, and s is the
slope of the linear regression δ13CPOC=f(T) in
‰ ∘C-1 obtained from Fig. 11a. There are significant
correlations between δ13C24∘C of biomass and
POC concentration (circles: R2=0.71, p<0.0001, n=18;
triangles: R2=0.66, p<0.0001, n=18; Fig. 11b), indicating that
primary production drives ∼ 70 % of the variation of phytoplankton
δ13C around DCM layers in the southern ECS. A similar positive
relationship between temperature-normalized δ13C and POC
concentration was observed by Savoye et al. (2003) during spring
phytoplankton blooms in the Bay of Seine, France. However,
δ13C24∘C correlated insignificantly with
the POC / Chl a ratio and C / N ratio (Fig. 11c and d), implying that
degradation has a minor effect on the carbon isotopic composition of POM in
this study.
Dynamics of δ15NPN in POM in DCM layers
In contrast to the POC and δ13CPOC relationship
(Fig. 10a), there is no significant relationship between PN and the isotopic
composition (δ15NPN) of the POM investigated in the
present study (Fig. 10b), implying that primary productivity has no
significant control on the variability of δ15NPN. As
the POM around the water depths of DCM was dominantly from the
newly produced, phytoplankton-synthesized source,
δ15NPN should be similar to δ15N in
phytoplankton. Considering the prevalence of low N / P ratios in the DCM
layer of the East China Sea (Lee et al., 2017), the degree of nitrate
utilization by phytoplankton should be high and that would result in the
composition of δ15NPN being similar to δ15N of
nitrate (δ15NNO3-) (Altabet and Francois, 1994;
Minagawa et al., 2001). Therefore, the spatial distribution of δ15NNO3- is probably crucial to decipher the distribution
of δ15NPN in DCM layers. Importantly, the spatial
distribution of δ15NPN (Fig. 5) resembles the surface
current pattern (Fig. 1), as well as the distribution of different water
masses (Fig. 7), suggesting that nitrate and the δ15NNO3- of CDW, TWCW and Kuroshio Water are largely
governing the distribution of δ15NPN in the study area.
According to Li et al. (2010), the range of δ15NNO3- in the Yangtze River was 7.3–12.9 ‰,
with a mean value of 8.3 ‰. In the north-east of Taiwan Island,
δ15NNO3- was 5.5–6.1 ‰ at depths of 500
to 780 m (Liu et al., 1996). However, TWCW is nutrient-depleted, enabling
incorporation of N-fixer-derived nitrogen in the suspended POM. This general
spatial pattern of δ15NNO3- (i.e. higher δ15NNO3- (> 6 ‰) in the north-east coastal
region and off north-east Taiwan, but lower δ15NPN in
between these two regions) exactly resembles the distribution of
δ15NPN in the DCM layers of this study (Fig. 5).
Therefore, the δ15NPN variation in the DCM layer of the
East China Sea was primarily governed by the nutrient status and
δ15NNO3-, though we do not have nutrient data
generated during the same cruise to validate our interpretations.
There is another possibility that high δ15NPN (DH7-8:
6.7 ‰, DH7-9: 7.8 ‰) in the DCM layer, off north-east Taiwan
(Fig. 5), may not result from the high degree of nitrate utilization, but
instead from the incorporation of inorganic nitrogen (mainly NH4+) in
the POM. According to Chen et al. (1996) and Liu et al. (1996), NO3-
and NH4+ concentrations in KSSW were high due to the decomposition of
OM in sinking particles. However, the concentrations of chlorophyll fluorescence as
well as POC and PN are low (Figs. 3 and 4). The low chlorophyll fluorescence might be
limited by the low temperature in this high-nutrient–low-chlorophyll region
(Umezawa et al., 2014). Because of the low temperature, the prevailing high
CO2 pressure is expected to decrease δ13C in DIC and may drive a
great carbon isotopic fractionation during carbon assimilation by
phytoplankton (Rau et al., 1992), the potential reason why
δ13CPOC values in these two stations were low (-25.8
and -25.2 ‰) compared to values of other locations in
the study area. Consistently, the low concentration of POC restricts the idea
that the high δ15NPN could not be from the
denitrification effect. The high δ15NPN (6.7,
7.8 ‰) values are probably due to the incorporation of inorganic nitrogen
(mainly NH4+); the process normally drives the δ15NPN as high as that of inorganic nitrogen δ15N
(Coffin and Cifuentes, 1999). Although δ15N of NH4+ in
Kuroshio Water is not available for comparison, it seems that δ15N
of remineralized NH4+ was relatively greater than δ15N of
NO3- (York et al., 2010). This possibility is also supported by the
high concentrations of NO3- and NH4+ in Kuroshio Subsurface
Water (Liu et al., 1996) as well as the low contents of POC (< 1; 0.96,
0.98 %) and low molar C / N ratios (4.1, 5.4) of these two SPM
samples (DH7-8 and DH7-9).
Impact of Yangtze River on POM in DCM of ECS
The range of POC / Chl a obtained in this study (33–200 g g-1)
is within the range (< 200 g g-1) reported for the
phytoplankton-dominated POM in the coastal and shelf waters (e.g. Chang et
al., 2003; Savoye et al., 2003; Hung et al., 2013; Liénart et al., 2016).
We also obtained a narrow range of C / N ratios (4.1–6.3), but a wide
range of δ13CPOC (-25.8 to -18.2 ‰)
compared to previous studies in the ECS (4.0–34.3, Liu et al., 1998; -24.0
to -19.8 ‰, Wu et al., 2003). Our results indicate that POM at the
DCM was largely produced in situ and derived from phytoplankton biomass,
with little terrestrial influence. The lack of terrestrial OM signals seems
to be related to reservoir- and dam-building along the river in recent years
that has shifted the location of the Yangtze-derived POC deposition from the
inner shelf of the ECS to terrestrial reservoirs (Li et al., 2015). The
sediment delivered from the river to the estuary has been reduced by 40 %
since 2003 when the Three Gorges Dam was completed (Yang et al., 2011,
and references therein). Recently, Dai et al. (2014) reported that the
particulate load discharged by the Yangtze has declined to
150 Mt yr-1, less than ∼ 70 % of its sediment delivery to
the ECS during the 1950s. Although 87 % of the mean annual sediment of
Yangtze River is discharged during the flood season from June to September
(Wang et al., 2007; Zhu et al., 2011), approximately 60 out of 87 % of
the fine-grained sediments are temporarily deposited near the estuary and
then later resuspended and transported southward along the inner shelf, off
mainland China (Chen et al., 2017, and references therein). The
Yangtze-transported POM moves up toward the north-east across the shelf along
the so-called Changjiang transport pathway in summer season (e.g. Gao et
al., 2014), which is largely affected by the combined effects of high river
discharge, south-west summer monsoon and the intensified TWC (Beardsley et
al., 1985; Ichikawa and Beardsley, 2002; Lee and Chao, 2003). The T–S
diagrams (Figs. 6 and 7) of this study also illustrate this view.
Accompanying the decreasing sediment input, dam-building in the Yangtze River
basin since 2003 has buried around 4.9 ± 1.9 Mt yr-1 biospheric
POC, approximately 10 % of the world riverine POC burial flux to the
oceans (Li et al., 2015). The POC flux from the Yangtze to the ECS (range:
1.27–8.5 × 1012 g C yr-1; Wang et al., 1989; Qi et
al., 2014) was significantly less than the estimated primary productivity
(72.5 × 1012 g C yr-1; Gong et al., 2003), implying the
predominance of marine-sourced organic matter in the ECS. Moreover, the
substantial quantity of organic substances that transported by the Yangtze
River may be completely modified before being ultimately deposited on the
inner shelf of the ECS and being transported further offshore (Katoh et al.,
2000; Lie et al., 2003; Chen et al., 2008; Isobe and Matsuno, 2008). Wu et
al. (2007b), for instance, observed an advanced stage of POM degradation in
the entire Yangtze River with an average degradation index of -1.1. Based
on the investigation of lipid biomarkers in a sediment core collected from
the ECS, Wang et al. (2016) suggested the dominant preservation of marine
autochthonous organic matter (∼ 90 %) in the ECS.