Temporal variability of stable carbon and nitrogen isotopic composition of size-fractionated particulate organic matter in the hypertrophic Sumida River Estuary of Tokyo Bay, Japan

Estuarine, Coastal and Shelf Science 68 (2006) 245e258 www.elsevier.com/locate/ecss

Temporal variability of stable carbon and nitrogen isotopic composition of size-fractionated particulate organic matter in the hypertrophic Sumida River Estuary of Tokyo Bay, Japan Taeko Sato 1, Toshihiro Miyajima*, Hiroshi Ogawa, Yu Umezawa 2, Isao Koike Marine Biogeochemistry Laboratory, Ocean Research Institute, The University of Tokyo, Minamidai 1-15-1, Nakano, Tokyo 164-8639, Japan Received 6 November 2005; accepted 11 February 2006 Available online 19 April 2006

Abstract To clarify the major factors that control stable carbon and nitrogen isotopic compositions (d13C and d15N) of suspended particulate organic matter in a hypertrophic estuary, seasonal variation in concentrations and stable isotopic compositions of particulate organic carbon and nitrogen (POC, PN), chlorophyll a content (Chl) and dry weight of suspended solid (SS) was studied in the Sumida River estuary of Tokyo Bay. Particulate material was fractioned into four size-fractions (3e20, 20e53, 53e250, 250e1000 mm). At the lower-estuarine (LE) site, the finer size fraction (<20 mm) was predominant (>80% Chl) during the autumnal bloom (AugeSep 2003). The middle size fractions (20e53 and 53e250 mm) however increased during the spring bloom (50e80% Chl; FebeApr 2004), reflecting difference in taxonomic composition of phytoplankton. At the upper-estuarine and the riverine sites, the <20 mm fraction was most abundant throughout the survey period. d13CPOC and d15NPN varied between
32 and
16& and between
8 and þ12&, respectively. Relationships of d13CPOC with POC/Chl and POC/PN ratios could be explained by assuming two end members: allochthonous (terrestrial) POC with a d13C close to
26.5&, and autochthonous (planktonic) POC whose d13C was variable and apparently correlated with d13C of ambient DIC (
10 to 0&) and phytoplankton abundance (Chl). d15NPN was apparently con13 15 trolled by water temperature and NHþ 4 concentration. The variability of d C and d N was different between size fractions: the apparent depen13 15 dence of d C on Chl as well as that of d N on water temperature was significantly stronger for the middle size fractions (mainly diatoms) than the finer fraction (mainly microflagellates), suggesting different physiological response between taxa. The second objective of this study was to evaluate contribution of terrestrial organic carbon (TPOC) to the POC size fractions by using d13C of isolated chlorophyll a at LE site as surrogate for the autochthonous d13C end member. During the autumnal bloom, TPOC was 23e36% on average for the middle size fractions (20e53 and 53e 250 mm), and 59% for the finest (3e20 mm) and coarsest (250e1000 mm) fractions. During the spring bloom, the average contribution of TPOC was generally low (21%) except for the finest fraction (36%). Negative correlation between salinity and calculated TPOC suggested that transport of terrestrial organic carbon to this estuary was controlled principally by the river discharge. Ó 2006 Elsevier Ltd. All rights reserved. Keywords: particulate organic matter; phytoplankton; size fractions; carbon isotope ratio; nitrogen isotope ratio; allochthonous organic carbon

1. Introduction In contrast to the open ocean in which most of organic carbon is present as dissolved organic carbon, suspended * Corresponding author. E-mail address: [email protected] (T. Miyajima). 1 Present address: Geo Information Department, ImageOne Co., Ltd. Shinjuku, Tokyo. 2 Present address: Department of Botany, University of Hawaii Manoa, Honolulu, HI, USA. 0272-7714/$ - see front matter Ó 2006 Elsevier Ltd. All rights reserved. doi:10.1016/j.ecss.2006.02.007

particulate organic carbon (POC) sometimes occurs in similar abundance to dissolved organic carbon in coastal and estuarine environments. Considering its rapid turnover rate, POC potentially has a great impact on ecological functions and biogeochemical properties of the land-sea interface zones (Heip et al., 1995; Turner and Millward, 2002). Estuarine POC comprises of several components of different origins, including plankton produced in situ, higher plant detritus transported by the river water, and resuspended organic sediment. In the case of estuaries that have developed and populated area in

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the watershed, inputs of domestic, agricultural and industrial runoff to the river water increase land-derived POC (primary pollution) on the one hand, and high concentrations of dis
solved nutrients such as NHþ 4 and NO3 in polluted river waters stimulate primary production in the estuaries, resulting in higher abundance of plankton-derived POC (secondary pollution) on the other hand. The increase of POC input causes enhanced oxygen consumption in seawater and sediment, and increased turbidity reduces light penetration and therefore the benthic primary production. Better understanding of the dynamics of POC is indispensable for effective management of estuarine ecosystems, especially of highly populated and industrialized areas. For quantitatively evaluating the relative contribution of land-derived (allochthonous) and plankton-derived (autochthonous) materials to POC and sediment organic carbon (SOC), stable carbon isotopic composition (d13C) has been conventionally used to discriminate between marine (highd13C) and terrestrial (low-d13C) organic carbon (Coffin et al., 1994). Assuming conservative mixing of terrestrial and marine organic matters, each with a more or less constant d13C, relative contribution of both sources can be estimated from d13C by isotope mass balancing. Similarly, stable nitrogen isotopic composition (d15N) often shows clear difference between terrestrial and marine origins (Peters et al., 1978; Mariotti et al., 1984; McClelland and Valiela, 1998; Umezawa et al., 2002), and has been used to evaluate terrestrial nitrogen input to estuaries in combination with d13C (e.g., Wada et al., 1987; Thornton and McManus, 1994; Middelburg and Nieuwenhuize, 1998). Interpretation of d13C and d15N signatures may not however be straightforward when the d13C and d15N values of terrestrial or marine end members are not always constant (Gearing et al., 1984; Ogawa et al., 1994). This factor is especially important for eutrophicated temperate estuaries, where seasonal blooms of phytoplankton are greatly enhanced and the isotopic signatures of the autochthonous POC and particulate nitrogen (PN) may fluctuate reflecting temporal change of metabolic activities (Cifuentes et al., 1988; Ogawa and Ogura, 1997; Hellings et al., 1999). Efforts have however rarely been made to evaluate possible temporal variability in the end member d13C and d15N values and finding their controlling mechanisms in previous studies of estuarine POC and PN. Tokyo Bay is one of the most eutrophicated basins of the world (Nomura, 1995, 1998). Red tide events and bottom hypoxia chronically occur in the inner area, with chlorophyll a concentration in the surface water often exceeding 100 mg l
1. Several previous studies have reported d13C and d15N of particulate and sedimentary organic matter in this bay (Wada et al., 1990; Ogawa et al., 1994; Ogawa and Ogura, 1997; Sukigara and Saino, 2005). Wada et al. (1990) suggested average terrestrial and marine end member d13C values in this bay of
26.5& and
20.3&, respectively, and d15N end members of þ1.8& and þ6.3&, respectively. Ogawa and Ogura (1997) elaborated spatial and seasonal variations of d13CPOC in several estuaries in this bay, and suggested that the variation of d13CPOC could be explained by a 3-end

member model, which comprised of marine POC (ca.
17&), terrestrial POC (ca.
27&), and estuarine POC (ca.
30&). Similarly low d13CPOC for estuarine end member has also been postulated in the case of the Schelde Estuary (Middelburg and Nieuwenhuize, 1998). All the previous studies in Tokyo Bay reported only d13C and d15N values of bulk (i.e., unfractionated) POC and PN. In this bay, the majority of plankton biomass usually consists of dinoflagellates and athecate microflagellates that belong mainly to the nanoplankton size range (2e20 mm), although larger organisms (>20 mm) such as chain-forming diatoms often predominate especially in spring (Nomura and Yoshida, 1997). The observed seasonal variations in d13C and d15N may have been partially caused by seasonal change in taxonomic composition of phytoplankton, which adds extra complexity to the interpretation of d13C and d15N (Gearing et al., 1984). To demonstrate internal variability of d13C and d15N within the suspended particle pool, Middelburg and Nieuwenhuize (1998) measured d13C and d15N values of size-fractionated particles of the Schelde Estuary and found a clear trend of the higher d13C and the lower d15N in the finer size fractions. The objective of the present study is to understand better seasonal and size-specific variability of d13CPOC and d15NPN and to improve the estimation accuracy for terrestrial contribution to the suspended particle pool using d13C signature in an estuary of Tokyo Bay. For this purpose, suspended particles were fractionated into defined size classes to facilitate assigning individual taxonomic groups of algae to specific size classes. d13C and d15N values for individual fractions were analyzed, and potential factors that affected d13CPOC and d15NPN of respective fractions were examined. As an attempt to follow seasonal change of the end member d13C, we measured d13C of isolated chlorophyll a to use as surrogate for the autochthonous-end member d13C (Qian et al., 1996; Otero et al., 2000) in quantitative evaluation of terrestrial POC input. Merits and possible problems of this approach are discussed, comparing with the traditional model that relies on static end member d13C values. 2. Materials and methods 2.1. Study site The Sumida River is a distributary of the Arakawa River, one of the major rivers inflowing to Tokyo Bay with the catchment area of ca. 2940 km2. The watershed of the Sumida River and its tributaries covers the most populated area of Tokyo City. Our survey was conducted twice a month from August 2003 to April 2004 at two stations of the Sumida River Estuary (Fig. 1): lower estuarine Station 1 was near the river mouth, and upper estuarine Station 2 was ca. 6 km upstream of Station 1 and ca. 22 km downstream of the separation point of the Sumida River from the Arakawa River. The range of tides at the river mouth is ca. 2 m, and water samples were collected within 3 h of the low tide. In addition, the same suite of sampling was done at another, riverine Station 3 (the Arakawa

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Fig. 1. Location of sampling stations along the Sumida River. Monthly average water levels (above the Arakawa Peil, which is approximately the low-water level of spring tide in Tokyo Bay) observed at Reiganjima (near Station 2) from January 2003 through December 2004 are also shown.

River, ca. 4 km upstream of the separation point of the Sumida River and ca. 29 km upstream of the Arakawa River mouth) three times (August 27 and December 9, 2003, and April 29, 2004). The hydrograph of the Sumida River (near Station 2) has been published on the WWW by Ministry of Land, Infrastructure and Transport Government of Japan (; available only in Japanese), and is summarized as monthly averages of water levels measured every hour in Fig. 1. 2.2. Sampling Surface water samples were collected from the piers using a clean plastic bucket. Water temperature was measured with a portable thermistor thermometer. Salinity was measured with a refractometer using the Practical Salinity Scale. Water samples for d13CDIC were subsampled into 30 ml serum vials with butyl rubber septa, without air bubbles being left inside, and immediately stabilized by adding 0.2 ml saturated HgCl2 solution. For the bulk particulate matter (PM), sample water was filtered at first through a 1 mm stainless mesh to remove floating large particles, and then PM was collected by vacuum filtration (<150 hPa pressure difference) on a preignited (450  C, 3 h) glassfiber filter (Whatman, GF/F; nominal pore size, 0.7 mm). A part of the filtrate was transferred to a 10 ml screw-capped acrylic tube for later measurement of inorganic nutrient concentration. Size fractionation of PM was carried out as follows. The sample water was treated

successively with 1 mm, 250 mm, and 53 mm stainless mesh screens and then a 20 mm nylon mesh screen. The volume of treated sample water varied by date depending on PM concentration of the water. Collected PM on each screen (except 1 mm) was washed back into a glass bottle with GF/F-filtered sample water. In addition, about 60 l of the sample water was directly filtered through the 20 mm nylon mesh, and collected PM was similarly treated as above and used later for isolation of chlorophyll a for isotopic analysis. The bulk and fractionated PM samples, nutrient samples, and 2e4 l sample water filtered through the 20 mm nylon mesh were transferred on ice to the laboratory. Then 20 mm-filtered sample water was filtered using a compressive filtration unit (ADVANTEC, UHP-43K) successively through a 3 mm- and 0.2 mm-mesh polycarbonate membranes (Nuclepore), and the retentates on the membranes were recovered in glass bottles. Each of the fractionated PM samples in glass bottles was collected on a preignited and preweighed glassfiber filter (GF/F) by vacuum filtration (<150 hPa pressure difference) for SS measurement and the elemental and isotopic analyses. A portion of the PM samples was collected on another preignited GF/F filter for chlorophyll a determination. The GF/F filters for elemental and isotopic analyses were immediately put in an drying oven (60  C), while those for chlorophyll a determination were soaked in 6 ml N,N0 -dimethylformamide and stored at
20  C in the dark. The size fraction <3 mm was analyzed only for chlorophyll a content. The GF/F filter

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containing >20 mm PM for chlorophyll isolation was placed in a brown glass vial that contained 50 ml of high-purity methanol þ acetone (7 þ 3), and the vial was stored at
85  C until further treatments. Water samples for nutrient analysis were stored at
20  C. Stabilized water samples for d13CDIC measurement were stored in the dark at room temperature. 2.3. Analyses Chlorophyll a concentration in the bulk and fractionated PM on GF/F was extracted by N,N0 -dimethylformamide and quantified using a fluorometer (Turner Design, 10-AU-005) (Suzuki and Ishimaru, 1990). The GF/F filters for SS determination and elemental and isotopic analyses were treated as follows. Filters were dried to a constant weight and the weight was recorded. The filters were then individually placed in Petri dishes, and the dishes were placed in a closed plastic container in which a cup of conc. HCl was also put. The filters were treated with HCl fumes overnight to remove inorganic carbon. The filters were then placed in a vacuum desiccator with some NaOH pellets for >12 h to completely remove acid vapor and CO2. Carbon and nitrogen contents, d13C and d15N of the filter samples were determined using an elemental analyzer connected online to an isotope-ratio mass spectrometer (Thermo-electron, FLASH EA-Conflo III-DELTAPLUS XP). Prior to analysis, the filters were folded and wrapped with tin capsules (prewashed with acetone) and molded into small tablets at a pressure of ca. 50 kgf cm
2. Isotopic compositions are expressed by the usual d-notation (Lajtha and Michener, 1994). Isolation of chlorophyll a from POM for isotopic analysis was carried out according to Otero et al. (2000). To 15 ml of the methanol þ acetone extract in a 50 ml polypropylene centrifuge bottle, 3 ml chloroform and 3 ml Milli-Q water were successively added and mixed well. The mixture was centrifuged at 4  C and 1500  g for 10 min. The bottom chloroform layer was transferred to a 30 ml glass vial and dried under a N2 stream at room temperature in a ventilated fume hood. The dried material was redissolved in 1 ml of NaHCO3-saturated acetone, and immediately injected to a preparative HPLC (Shimadzu, LC-6AD). The HPLC was equipped with a reverse-phase column (Shiseido, CAPCELL PAK C18, 5 mm grain size, 20 mm diameter, 250 mm length), a photodiode array detector (Shimadzu, SPD-M10AVP) and a fraction collector (Shimadzu, FRC-10A, with a sample cooler at 4  C). Chlorophylls in the sample were eluted by methanol/ acetone gradient at a flow rate of 20 ml min
1, and the chlorophyll a fraction, as determined by the absorbance spectrum, was collected into preignited glass vials. The collected fraction was pooled and dried up under a N2 stream at 40  C. The dried residue was redissolved with 0.5 ml dichloromethane, and the solution was dropped into a tin cup for elemental analysis of liquid samples (precleaned with acetone) using a preignited Pasteur pipette, and dried up again under a N2 stream at room temperature. The tin cups containing isolated chlorophyll a were analyzed using a mass spectrometer (Thermo-electron, FLASH EA-Conflo III-DELTAPLUS XP).

In principle, both d13C and d15N of chlorophylls can be determined by this protocol; however, our isolated chlorophyll samples were often so small that the obtained d15N values were unreliable. Therefore, we report only d13C value for isolated chlorophyll samples in this study. d13C of DIC was measured using the headspace method according to Miyajima et al. (1995). Briefly, a helium headspace (5 ml) was created within the glass vial containing Hg-stabilized sample water, and then the content was acidified with 0.3 ml of degassed 6.0 M HCl, and equilibrated by vigorous hand-shaking. A 0.1e0.2 ml portion of the headspace gas was manually injected into a GC-C-IRMS system (Thermoelectron, GC6890-Combustion III-DELTAPLUS XP). The GC was equipped with a capillary column (J&W, GS-GAS-PRO, 30 m length, 320 mm inner diameter) and operated with oven and injector temperature of 70  C, He carrier flow rate of 1.5 ml min
1 and a split ratio of 20:1. d13C of DIC was calculated from the measured d13C of the headspace CO2 using appropriate solubility constant and equilibrium isotope fractionation factor (Mook et al., 1974). Concentrations of dissolved inorganic nutrients (NHþ 4,
3
NO
2 , NO3 , PO4 ) were determined using a nutrient analyzer (BRAN þ LUEBBE, AACS-III). Statistical analysis and curve fitting were carried out using software packages StatView ver. 5 (SAS Institute) and pro Fit ver. 6 (QuantumSoft), respectively. 3. Results Water temperature varied between 9.5  C (Jan 26) and 27.0  C (Sep 18), with little difference (1.1  C) between lower estuarine Station 1 and upper estuarine Station 2 (Fig. 2). Salinity was always higher at Station 1 than Station 2. Higher salinity in winter than in autumn and spring corresponded partly to reduced discharge of the Sumida River in winter (Fig. 1) due to poorer precipitation in the watershed, and partly to enhanced vertical mixing of water column due

to cooling of surface water. Nutrients (NHþ 4 , NO3 , NO2 , 3
PO4 ) were always much higher at Station 2 than Station 1. The atomic ratio of dissolved inorganic nitrogen (DIN) to phosphorus (DIP) was always considerably higher than the Redfield ratio (28e151). At riverine Station 3, water temperature was similar to the other stations, but salinity was always virtually
0. NO
3 (60e90 mM) and NO2 (0e12 mM) were in similar ranges to those at Station 1. NHþ 4 was however higher (130e 560 mM) and PO3
was lower (0.4e1.0 mM) than the down4 stream stations. DIN/DIP ratio was higher than 200. Concentration of NO
3 was significantly negatively correlated with salinity (Fig. 3), indicating that the riverine input þ is the predominant source of NO
3 in this estuary. For NH4 3
and PO4 , negative correlation with salinity was statistically significant but the linearity was not so clear as the case of 13 NO
3 . d CDIC showed clear positive correlation with salinity, from which d13C of river water DIC at salinity ¼ 0 was expected to be
13.1&. Concentrations of POC and PN were clearly higher in autumnal (AugeSep) and spring (FebeApr) blooming periods

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3
Fig. 2. Seasonal changes in water temperature, salinity, and nutrient concentrations (NHþ 4 , NO3 , NO2 , PO4 ) during the survey period (August 2003eApril 2004) at Station 1 (solid circle) and Station 2 (open circle).

than non-blooming period (OcteJan) (Fig. 4). They were also generally higher at Station 1 than Station 2 in the blooming periods while the reverse was the case in the non-blooming period. Size distribution of POC and PN was different between the two blooming periods. At Station 1, POC and PN of 3e20 mm size fraction were dominant (>70%) during the autumnal bloom, while those of coarser fractions (>20 mm) increased to 30e57% altogether during the spring bloom. At Station 2, 3e20 mm fraction usually dominated POC and PN pools, although fraction of >20 mm particles temporarily increased in February. Bulk POC/PN atomic ratio was usually 6e7 at Stations 1 and 3, and 7e10 at Station 2. POC/PN ratio was often considerably higher in 250e1000 mm fraction than in the other fractions, especially at Stations 2 and 3. Abundance and size distribution of Chl (Fig. 5) showed similar seasonal pattern to POC and PN. The difference in

size distribution between the two bloom periods was even more conspicuous in Chl than POC and PN. Concentration of Chl was usually higher at Station 1 than Station 2. Bulk POC/Chl weight ratio, which is an indicator of the relative abundance of aged detrital POC to fresh phytoplanktonic POC, ranged between 50e1000 at Stations 1 and 2, and around 1000 at Station 3. POC/Chl ratio was high (up to 20,000) in 250e1000 mm fraction as compared with the other fractions, especially at Station 1. Seasonal change in SS (Fig. 5) was poorly related to those of POC, PN or Chl, although the size distribution pattern was similar between PN and SS. Percentage of POC in total SS of the bulk particulate matter was usually 5e20 wt.% at Stations 1 and 2, and >20 wt.% at Station 3. Similarly to POC/Chl ratio, relative portion of POC in total SS in 250e1000 mm fraction was usually higher than in the other fractions at Stations 1 and 2.

13 þ 3
Fig. 3. Correlation of concentrations of NO
3 , NH4 and PO4 and d C of dissolved inorganic carbon to salinity. Data from Station 1 (solid circle) and Station 2 (open circle).

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Fig. 4. Seasonal changes in concentrations of POC and PN and their size fractions at three stations from August 2003 to April 2004. Bulk concentration (open circle) and the sum of the four size fractions (solid circle) were shown for POC and PN. Color gradient from thick to thin gray corresponds to size fractions from small to large: 3e20, 20e53, 53e250, and 250e1000 mm.

Seasonal variations of d13CPOC and d15NPN of the size fractions are shown in Fig. 6. The magnitude of seasonal variation was apparently larger at Station 1 than Stations 2 and 3. Difference between size fractions was usually within 10& for both d13C and d15N. The bulk d13CPOC and d15NPN were sometimes lower than those of all the size fractions between the end of September and the beginning of January (Fig. 6). This indicates that particulate matter (PM) of <3 mm, which occupied a significant fraction of the PM pool in this period as illustrated by the size distribution of Chl (Fig. 5), had considerably lower d13C and d15N values than the coarser fractions (3e1000 mm). d13C of chlorophyll a (d13Cchl) isolated from >20 mm PM at Station 1 (Fig. 6) was usually similar to d13CPOC of >20 mm fraction. For sampling dates on which d13Cchl could be measured, Chl concentration in the >20 mm fraction was 6.3 times on average as high at Station 1 as at Station 2, which indicates that the measured d13Cchl represented that of phytoplankton that had grown downstream below Station 2. 4. Discussion 4.1. Characteristics of d13C and d15N dynamics in POM size fractions The difference in size distribution of Chl, POC and PN between the autumnal and spring bloom periods (Figs. 4, 5;

especially at Station 1) can be attributed to seasonal change in taxonomic composition in phytoplankton community. Phytoplankton community in the inner region of Tokyo Bay consists mainly of chain-forming centric diatoms (such as Skeletonema costatum, Chaetoceros spp., Thalassiosira spp.), dinoflagellates (such as Prorocentrum spp., Gymnodinium spp.) and small athecate flagellates (such as Heterosigma akashiwo) (Nomura and Yoshida, 1997; Nomura, 1998). Small flagellates and dinoflagellates usually predominate phytoplankton biomass (up to 100 mg Chl l
1) in warmer seasons of the year. Diatoms often increase their relative abundance to >1000 cells ml
1 in spring, and sometimes form shortterm blooms also in summer. Although the individual cell size of diatoms belongs to both 3e20 and 20e53 mm size fractions, they usually form long chains of several tens of cells, and thus mainly occur in > 20 mm size fractions. Thus, the gradual increase of relative abundance of Chl in 20e53 and 53e250 mm fractions from December to March (Fig. 5) can be attributed to growth of diatom populations. In contrast, the predominance of 3e20 mm fraction of Chl in August and September indicates that the bloom of this season consisted almost exclusively of small microflagellates. The d13CPOC and d15NPN in the estuary is principally determined by d13C and d15N of terrestrial PM delivered by the river and autochthonously produced PM derived mainly from phytoplankton, and the mixing ratio of these two pools. Original

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251

Fig. 5. Seasonal changes in concentrations of chlorophyll a and suspended solid (SS) at three stations from August 2003 to April 2004. Symbols and colors are same as in Fig. 4. For chlorophyll a, separately determined <3 mm size fraction is indicated by the thickest gray, and the plot of the bulk concentration is omitted.

terrestrial plant remains may be highly diverse in d13C (Michener and Schell, 1994), including C3 plants (usually around
27&), C4 plants (around
14&), submerged C3 plants and attached algae (
30 to
10&). These source organic matters are mixed and averaged during decomposition, erosion and

transport processes, and as a result, terrestrial PM in estuaries takes relatively constant d13C between
26 and
31& (Coffin et al., 1994). If d13CPOC and d15NPN in the estuary had been determined simply by conservative mixing of static terrestrial (low d13C, low d15N) and marine (high d13C, high d15N) end

Fig. 6. Seasonal changes in d13CPOC and d15NPN of the bulk and the size fractions of particulate matter at three stations, and d13C of chlorophyll a extracted from >20 mm particulate matter collected at Station 1 from August 2003 to April 2004. Solid circle, 3e20 mm fraction; solid triangle, 20e53 mm fraction; open triangle, 53e250 mm fraction; open circle, 250e1000 mm fraction; gray diamond, chlorophyll a; broken line, the bulk particulate matter.

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252

members, they would have covaried keeping positive correlation to each other (Wada et al., 1990). In our case, d13CPOC was higher in the order: spring bloom > winter non-blooming period > autumnal bloom. However, d15NPN varied in the reverse manner (Table 1). This trend is not consistent with the conservative mixing model. Furthermore, C:N and POC:Chl ratios of PM, which have been used as indicators of the ratio of terrestrial to autochthonous organic carbon in estuarine POC (Thornton and McManus, 1994; Hellings et al., 1999), showed no correlation to d13CPOC (Fig. 7). d13CPOC varied widely from
32 to
16& when POC:Chl and C:N were relatively low, that is, when PM consisted mainly of autochthonous particles. As POC:Chl and C:N increased beyond 200 and 10, respectively, d13C converged irrespectively of stations toward a narrow range around
26.5&, which is the terrestrial end member d13C value postulated by Wada et al. (1990). Similarly, d15NPN varied from
8 to þ10& when C/N <12 and POC/ Chl <1000, while it gradually converged to between þ1 and þ4& as C/N and POC/Chl became higher. Similar trends have been reported for the Delaware Estuary by Cifuentes et al. (1988). These facts imply that the major factor that affected seasonal variations in d13CPOC and d15NPN in the estuary is internal isotopic variability of the autochthonously produced PM such as phytoplankton, rather than the mixing ratio of terrestrial vs. autochthonous PM. The following sections examine dependence of d13CPOC and d15NPN on environmental and physiological factors that would primarily affect isotopic composition of autochthonous PM. 4.1.1. Factors controlling d13C d13CPOC of phytoplankton is determined by d13C of substrate DIC and the magnitude of isotope discrimination during photosynthesis. DIC of seawater takes relatively constant d13C between 0 and þ2&, while d13C of riverine DIC is usually lower than 0&; as a result, d13C of estuarine DIC is determined principally by mixing ratio of marine and freshwater DIC, showing strong correlation to salinity (Fig. 3). The

magnitude of isotope discrimination during photosynthesis, on the other hand, depends on growth rate and pCO2 (Goericke et al., 1994; Laws et al., 1995; Popp et al., 1998; Riebesell et al., 2000) as well as taxonomic group of phytoplankton (Wong and Sackett, 1978; Hinga et al., 1994; Burkhardt et al., 1999). In general, higher growth rate leads to smaller discrimination, and consequently higher d13CPOC. As water temperature increases, pCO2 and isotope discrimination between HCO
3 and CO2(aq) decrease and algal growth rate increases, which altogether results in apparent positive correlation between d13C of phytoplankton and water temperature (Wong and Sackett, 1978; Fontugne and Duplessy, 1981; Savoye et al., 2003). As average d13C of marine phytoplankton increases with water temperature at ca. 0.2e0.5&  C
1 (as reviewed by Savoye et al., 2003), the variation of water temperature of ca. 17  C during the study period (Fig. 2) could have caused a variation of about 3.4e8.5& for d13C of phytoplankton. On the other hand, changes of salinity were 19 and 28 at Stations 1 and 2, respectively, which correspond to changes of d13CDIC of ca. 6 and 9&, respectively (Fig. 3). Thus, the potential effects of temperature on d13CPOC should have been of similar magnitude to that of d13CDIC. Hence, dependences of d13CPOC on salinity, temperature and Chl concentration were examined for Stations 1 and 2. For the bulk POC as well as all the size fractions, d13CPOC was strongly correlated with salinity (r ¼ 0.500e0.737; p < 0.002). As salinity does not directly influence isotope discrimination during photosynthesis (Wong and Sackett, 1978), the dependence of d13CPOC on salinity suggests that d13C of phytoplankton was controlled by d13CDIC in the surrounding water. Significant correlation of d13CPOC to temperature ( p < 0.04) was also detected for the bulk POC and all the size fractions. Correlation to temperature was however always negative (r ¼
0.349 to
0.481) and thus opposite to the expected relationship. This is also true when the correlation is evaluated separately for the respective Stations. This apparent temperature dependence was presumably brought about as a result of negative correlation between water temperature and salinity (r ¼
0.517, n ¼ 36; Fig. 2), and implies that the proper

Table 1 d13CPOC and d15NPN of the bulk particulate matter and the size fractions, as averaged over the autumnal blooming period (Period 1, AugeSep 2003), the winter non-blooming period (Period 2, Oct 2003eJan 2004) and the spring blooming period (Period 3, FebeApr 2004) d13C (&)

d15N (&)

Period 1

Period 2

Period 3

Period 1

Period 2

Period 3

Station 1 Bulk 250e1000 mm 53e250 mm 20e53 mm 3e20 mm


25.6  1.8
24.8  1.2
23.2  0.8
23.7  1.5
25.3  1.8


26.0  1.8
24.1  1.5
22.5  1.7
22.7  1.4
25.3  1.1


20.0  1.4
19.1  2.1
18.9  1.6
19.5  1.3
20.3  2.8

2.1  5.5 5.1  2.0 7.5  1.4 6.0  1.7 2.8  3.2

1.2  1.7 5.9  3.0 5.9  3.4 3.8  2.6 2.4  0.5


1.6  0.8 1.7  4.4 1.0  2.8 0.3  2.5
0.6  1.3

Station 2 Bulk 250e1000 mm 53e250 mm 20e53 mm 3e20 mm


27.4  1.2
24.2  1.3
26.0  2.4
26.3  1.3
28.6  2.2


25.8  1.2
24.9  1.7
24.5  0.6
25.0  1.0
25.6  1.3


24.7  1.5
23.5  3.0
23.5  2.8
24.6  1.1
25.6  2.0

0.7  2.2 4.8  0.9 5.2  0.9 4.2  1.1 1.9  2.4

0.7  1.9 3.1  2.3 3.8  2.3 2.8  1.8 1.4  1.0


1.6  2.2 2.2  2.4
0.8  4.8
0.8  2.5
0.3  1.4

T. Sato et al. / Estuarine, Coastal and Shelf Science 68 (2006) 245e258

253

Fig. 7. d13CPOC vs. POC/PN atomic ratio (left) and POC/Chl weight ratio (right) for the bulk POM and four size fractions. Hatched areas A and B indicate d13C end members for allochthonous (terrestrial) and autochthonous (planktonic) POM, respectively, as expected from the data.

effect of water temperature on phytoplankton d13C was relatively minor as compared to the effect of d13CDIC in our study site. The growth rate of phytoplankton in such a hypertrophic environment as our study site does not simply depend on water temperature, but may be affected even more strongly by other factors such as nutrient conditions, water mixing intensity, and taxonomic composition of phytoplankton. This is presumably a reason why the effect of temperature on d13CPOC was apparently so weak in our study site. In contrast, clearly higher and lower d13CPOC in summer and winter, respectively, have been reported for the offshore area and the bay mouth of Tokyo Bay (Ogawa and Ogura, 1997; Sukigara and Saino, 2005), where freshwater DIC and nutrient inputs are much lower than our study site. Since growth rate is defined as time-derivative of the logarithm of biomass (Chl), seasonal change in logarithmic Chl is expected to reflect seasonal change in growth rate of phytoplankton to some extent. Correlation of d13CPOC to logarithmic Chl was always positive, and statistically significant (r ¼ 0.389e0.580; p < 0.02) except for the <20 mm fraction. This result suggests that the effect of growth rate on

d13CPOC was also remarkable. In addition to the direct effect of growth rate on carbon isotope fractionation, high phytoplankton density as indicated by high Chl should have resulted in low pCO2, which in turn might have forced phytoplankton to use bicarbonate instead of dissolved CO2 for photosynthesis. This is another possible mechanism for the positive correlation between d13CPOC and Chl. Dependence of d13CPOC on salinity and logarithmic Chl was evaluated by the multiple regression analysis (MRA, Table 2) to compare influence of d13CDIC and algal density. Correlations to salinity and Chl were equally strong for the size fractions >20 mm. However, dependence on Chl diminished for the 3e20 mm fraction. For the bulk POC, significant correlation of d13CPOC was found with both salinity and Chl, but correlation with Chl was much weaker than that with salinity. These results suggest that, although dependence of cellular d13C on substrate d13CDIC generally applies to all the size fractions of phytoplankton, dependence of carbon isotope discrimination on growth rate or density is much stronger for coarser size fractions of phytoplankton, such as diatoms, than for

Table 2 Standard regression coefficients and r2 values of multiple regression analyses of d13C values to salinity and logarithmic chlorophyll a concentration (left columns) and of d15N values to water temperature and NHþ 4 concentration (right columns) for the bulk and size-fractionated particulate matter. Data of Stations 1 and 2 are compiled (n ¼ 36). *, ** and *** indicate that the regression is statistically significant at p < 0.05, p < 0.01 and p < 0.001, respectively d13CPOC

Bulk POM 250e1000 mm fraction 53e250 mm fraction 20e53 mm fraction 3e20 mm fraction

d15NPN
1

r

Salinity

Ln[Chl (mg l )]

r2

TW (  C)

NHþ 4 (mM)

0.447 0.480 0.498 0.696 0.468

0.545*** 0.420** 0.430** 0.630*** 0.662***

0.353* 0.487*** 0.429** 0.405*** 0.225

0.298 0.154 0.516 0.471 0.271

0.431** 0.214 0.557*** 0.540*** 0.313


0.209
0.276
0.309*
0.300*
0.278

2

254

T. Sato et al. / Estuarine, Coastal and Shelf Science 68 (2006) 245e258

smaller phytoplankton such as microflagellates. The average d13CPOC of the finest 3e20 mm fraction that was slightly more negative than the coarser fractions (Table 1) may also be related to taxonomic or cell-size-specific difference in carbon isotopic discrimination during photosynthesis (cf. Popp et al., 1998; Burkhardt et al., 1999). These differences may be explained alternatively by different contribution of terrestrial organic carbon between size fractions. As discussed later, more than half of POC in the 3e20 mm fraction during the autumnal bloom and the winter non-blooming period might have been of terrestrial origin, and consequently the isotopic signal of autochthonous POC might have been considerably diluted. Thus, further studies are necessary to identify the mechanisms underlying these differences between size classes from both physiological and hydrographic points of view. 4.1.2. Factors controlling d15N d15NPN generally decreased with time for both the bulk PM and the size fractions, and the difference between the autumnal and spring blooms was larger for 53e250 mm and 20e53 mm fractions than the other fractions for both stations (Fig. 6, Table 1). d15NPN of autochthonously produced PM is principally determined by d15N values of substrate N, isotope fractionation factor (3) during N uptake by phytoplankton, and the fractional utilization of substrate N (Sigman et al., 1999; Altabet and Francois, 2001). In contrast to carbon, the substrate N may consist of several chemical species including

NHþ 4 , NO3 , NO2 , urea, amino acids, and also N2 in the case of N2-fixing algae. In our study site where concentrations
of NHþ 4 and NO3 were extremely high (Fig. 2), it can be assumed that these two species were virtually the only N source for phytoplankton. Unfortunately, we do not have d15N values of either of these species. If d15N was different between NHþ 4 15 and NO
3 , then d N of phytoplankton should be affected by
relative availability of NHþ 4 to NO3 . The 3 value of nitrogen þ uptake depends on substrate (NH4 vs. NO
3 ) (Pennock et al., 1996; Waser et al., 1998a), growth conditions such as DIN concentration and light (Waser et al., 1998b; Needoba et al., 2004), and taxonomic group of phytoplankton (Montoya and McCarthy, 1995; Needoba et al., 2003). Given an 3 value, d15N of phytoplankton becomes higher as the fraction of utilization of substrate N becomes larger (i.e. the supply of DIN becomes depleted relatively to the demand), which would result in negative correlation between d15N of phytoplankton and DIN concentration. Significant negative correlation between d15NPN and NHþ 4 concentration was found for all the size fractions (r ¼
0.440 to
0.331; p < 0.05) except for the bulk PM. This is naturally ascribed to isotope discrimination associated with the degree of NHþ 4 utilization as mentioned above. However, d15NPN showed no correlation to NO
3 concentration ( p > 0.2, r2 < 0.05). This suggests that NO
3 was not an important nitrogen source for phytoplankton in this estuary. This interpretation is consistent with the facts that NHþ 4 , which is preferentially incorporated to NO
by most phytoplankton, 3 was never depleted from the water column of this estuary þ (Fig. 2). Stronger correlation of NO
3 than NH4 with salinity

(Fig. 3) also implies conservative behavior of NO
3 in this estuary. The apparent dependence of d15NPN on NHþ 4 but not on NO
3 have also been reported even in an offshore site
of inner Tokyo Bay, where both NHþ 4 and NO3 concentrations are much lower than at our study site (Sukigara and Saino, 2005). On the other hand, d15NPN was significantly positively correlated with temperature for the bulk PM and the size fractions (r ¼ 0.385 to 0.637; p < 0.02) except for 250e1000 mm. In contrast to the case of d13C, correlation between d15NPN and logarithmic Chl concentration was always negative and, except for the 53e250 mm fraction, statistically insignificant ( p > 0.07, r2 < 0.1), which implies that growth rate may not have been one of the major factors that affected d15NPN. Although we do not have direct evidence, dependence of d15NPN on temperature may be ascribed also to seasonal variation in d15N of NHþ 4 (Cifuentes et al., 1989). Due to the kinetic isotope fractionation by evaporation of NH3 and algal uptake of NHþ 4 , which are both enhanced by high water temperature, d15N of NHþ 4 remaining in the estuarine water might have elevated in warmer seasons. d15N of phytoplankton that grew on remaining NHþ 4 should have also increased, resulting in apparent dependence of d15NPN on water temperature. In addition, NHþ 4 could also be supplied from bottom water of the bay by vertical mixing especially in winter. This seasonal input of bottom water NHþ 4 may have been an additional cause for the seasonal variation of d15NPN, if it had a different d15N signature from the riverine NHþ 4. We also conducted MRA to compare the effects of water 15 temperature and NHþ 4 concentration on d NPN (Table 2). The effect of temperature was larger than that of NHþ 4 for all the size fractions as well as for the bulk PM. If our above interpretation for temperature dependence of d15NPN is correct, this result indicates that variation of d15N of autochthonously produced PM was principally controlled by isotope fractionation processes that operated on NHþ 4 preceding the production of PM, rather than isotope fractionation associated with the production of PM. However, this point should also be thoroughly examined in future studies. MRA also showed that the temperature effect on d15N was more conspicuous for 20e 53 and 53e250 mm fractions than the <20 mm fraction and the bulk PM, implying some physiological difference between diatoms and microflagellates. 4.2. Estimation of terrestrial contribution to estuarine POC pool The fact that average terrestrial PM shows consistent difference from PM produced by pelagic phytoplankton in d13CPOC (and sometimes also d15NPN, C/N ratio, POC/Chl ratio) has been used for discriminating the origins of suspended and sedimentary PM in estuarine and coastal environments (Peters et al., 1978; Wada et al., 1987, 1990; Cifuentes et al., 1988; Thornton and McManus, 1994; Raymond and Bauer, 2001; Usui et al., 2006). However, actual estuarine d13CPOC and d15NPN are often significantly deviated from the values predicted by conservative mixing of terrestrial and marine end

T. Sato et al. / Estuarine, Coastal and Shelf Science 68 (2006) 245e258

members, because these values of newly produced PM in estuarine water are strongly affected by biological factors such as growth rate, species composition and food web activities (Owens, 1985; Cifuentes et al., 1988; Sigleo and Macko, 2002; see also Section 4.1). This fact hampers us from quantitative estimation of terrestrial contribution to estuarine PM pool using simple mixing model that relies on static end member d13CPOC and d15NPN values. On the other hand, it has been demonstrated that the photosynthetic pigment chlorophyll a isolated from planktonic algae or terrestrial plants usually takes very similar d13C to the whole plant material (Bidigare et al., 1991; Kennicutt et al., 1992; Laws et al., 1995; Sachs et al., 1999; Otero et al., 2000). In fact, d13C of chlorophyll a (d13Cchl) isolated from estuarine PM has been used as surrogate for d13C of autochthonously produced POC (Qian et al., 1996; Otero et al., 2000). Our objective in this section is to evaluate to what extent the estimates of the fraction of terrestrial organic matter in the estuarine POC pool at Station 1 by carbon isotope mass balancing may be improved by using d13Cchl as surrogate for autochthonous d13CPOC (dynamic end member model, Model I), as compared to the estimates by assuming a fixed end member d13C (
20.3&; Wada et al., 1990) (static end member model, Model II). For terrestrial (allochthonous) POC, the bulk d13CPOC at the upstream Station 2 and a fixed end member d13C of
26.5& (Wada et al., 1990) were used for Models I and II, respectively. Fractions of terrestrial POC to the bulk and size-fractionated POC pools were estimated by these two approaches for each of the sampling dates on which d13Cchl was successfully determined, and then averaged over the two bloom periods and the non-blooming period (Table 3). In general, the fraction of terrestrial POC was larger for the autumnal bloom than for the spring bloom. This trend is consistent with seasonal change of discharge of the Sumida River (Fig. 1). The fraction of terrestrial POC was highest in the finest (3e20 mm) fraction and relatively low in the middle fractions (20e53, 53e250 mm). In the autumnal bloom, the estimated fraction of terrestrial POC for the bulk POC was even higher than that for 3e20 mm fraction. Considering that Table 3 Fraction (%) of terrestrial organic carbon in bulk and size-fractionated POC, as estimated from d13C values using two different models Perioda

Model

Bulk POM

250e 1000 mm

53e 250 mm

20e 53 mm

3e 20 mm

Period 1 (AugeSep 2003; n ¼ 3) Period 2 (Oct 2003eJan 2004; n ¼ 3) Period 3 (FebeApr 2004; n ¼ 6)

Ib IIc

67.4 81.1

52.8 75.0

23.7 52.6

36.4 59.1

59.5 75.7

I II

73.3 73.4

31.2 46.2

18.3 24.7

10.5 17.8

73.2 75.3

I II

31.3 6.7

20.9 5.0

13.6 3.8

20.0 3.0

35.9 18.0

a

Definition of periods is the same as in Table 1. Model I uses d13CChl as marine end member and d13C of the bulk POC as Station 2 as the terrestrial one for respective sampling dates. c Model II uses static end member values:
20.3& for marine and
26.5& for terrestrial (Wada et al., 1990). b

255

the bulk POC included also <3 mm fraction of POC, this result implies that <3 mm fraction consisted mostly of terrestrial POC in this season. The relatively larger terrestrial contribution to the bulk POC and 3e20 mm fraction suggests that terrestrial POC transported by the Sumida River was confined in relatively fine-size fraction compared to phytoplankton that grew within the estuary. On the other hand, the estimated fraction of terrestrial POC was also higher in 250e1000 mm fraction than 20e53 and 53e250 mm fractions for all the three periods. This is partly because phytoplankton that belong to such a large-size fraction are relatively rare in the study site. The 250e1000 mm fraction was always only a small fraction of the bulk POC (Fig. 4). In contrast, the <20 mm fraction usually occupied more than half of the bulk POC. These facts mean that terrestrial organic matter transported by the Sumida River mostly consisted of organic matter associated with siltand clay-sized particles, which was presumably originated from erosion of organic soils in the watershed. Difference in estimated fraction of terrestrial POC between the autumnal and spring blooms by Model II (4.2e19.4 times) was considerably larger than that by Model I (1.6e2.5 times) (Table 3). Considering that POC and Chl concentrations at Station 1 were not so much different between the autumnal and spring periods (Figs. 4, 5), such a big difference in fraction of terrestrial POC estimated by Model II is apparently an artifact. Furthermore, mass balance calculation with Model II often yielded erroneous estimates for fraction of terrestrial POC, such as <0% and >100% (22 and 3 out of 60 cases, respectively). These erroneous values were treated as 0 and 100%, respectively, when calculating the average values in Table 3. In contrast, calculation using Model I yielded such erroneous estimates only in 5 and 2 out of 60 cases, respectively. This fact also shows that Model I gives more realistic and accurate estimates than Model II. One problem associated with Model I is possible fluctuation in d13C of isolated chlorophyll a from the whole phytoplankton cell. Sachs et al. (1999) argued that the difference of d13C of the whole cell relative to chlorophyll a (Dd13CcellChl) is likely to be between
1.86 and þ0.15&, though it sometimes deviated beyond this range in incubated algal cells, presumably due to variability of growth rate that enhances isotopic heterogeneity within the chlorophyll molecule. In our case, if Dd13Ccell-Chl of
1.86 and þ0.15& is assumed, then the fraction of terrestrial POC in the bulk POC is estimated to be 52.6 and 68.5%, respectively, for the autumnal bloom and 10.8 and 32.8%, respectively, for the spring bloom; thus, the maximal uncertainty of 15.9e22.0% is expected for terrestrial contribution to POC size fractions. This limit of uncertainty theoretically increases when the difference between d13Cchl and the terrestrial end member d13C becomes small, and/or when autochthonous POC predominates the POC pool. Finally, we calculated concentration of terrestrially derived POC at Station 1 by multiplying the fraction of terrestrial POC estimated by Model I by the concentration of total POC, and plotting them against salinity for the bulk POC (Fig. 8). For comparison, concentration of total POC is also shown. Correlation with salinity was statistically significant for the calculated terrestrial

256

T. Sato et al. / Estuarine, Coastal and Shelf Science 68 (2006) 245e258

Fig. 8. Total POC (open circle) and terrestrially derived POC as calculated by Model I (solid circle) are plotted against salinity. Curve fitting was applied to terrestrial POC using a hyperbolic formula: x ¼
a $ y þ b/y þ c, where x is salinity, y is POC, a, b and c are coefficients to be fitted, assuming that errors for salinity measurement and the estimation by Model I are 0.5 (as Practical Salinity Scale) and 10% of total POC, respectively.

POC ( p ¼ 0.022) but not for total POC ( p ¼ 0.076). This correlation suggests that concentration of terrestrial POC was largely determined by conservative mixing of the Sumida River water with seawater. However, the relationship between concentration of terrestrial POC and salinity was represented well by a hyperbolic curve rather than a linear regression line (Fig. 8). This fact may be explained by two hypotheses: (i) retention of terrestrial POC in the water column was limited by salt-induced coagulation and sedimentation of fine organic particles, especially when salinity exceeded 25; and/or (ii) concentration of terrestrial POC in the river water was somewhat reduced when the river discharge rate was low. Correlation between salinity and estimated terrestrial POC was generally weak for the size fractions >20 mm of POC, however. The validity of these hypotheses remains to be examined in future studies. 5. Conclusions This study has described seasonal changes in d13C and d15N of four size fractions of PM, and evaluated the utility of d13C signature of chlorophyll a (d13Cchl) for quantitative estimation of land-derived and autochthonous organic carbon in estuarine POC pool in the hypertrophic Sumida River Estuary of Tokyo Bay. Major findings were: (1) average d13CPOC was higher in the microplankton fractions (20e53 and 53e250 mm; mainly diatoms) than in the nanoplankton fraction (3e20 mm; mainly microflagellates) and the mesoplankton fraction (250e 1000 mm). d15NPN was generally the higher in the coarser size fractions. (2) d13CPOC widely varied from
32 to
16& when the PM pool was dominated by planktonic organic matter

(low POC/Chl), whereas it converged to a narrow range around
26.5& when the PM pool was dominated by detrital particles (POC/Chl >200). d15NPN also showed a similar trend. This indicated that the variability of estuarine d13CPOC and d15NPN could be accounted for by the mixing model between terrestrial PM with relatively constant d13CPOC and d15NPN and autochthonous (planktonic) PM with highly variable d13CPOC and d15NPN. (3) d13CPOC in all the size fractions was controlled by d13C of the substrate DIC in estuarine water. The effect of algal density or growth rate on d13CPOC was also suggested for the meso- and microplankton fractions but not for the nanoplankton fractions. Water temperature barely affected d13CPOC. (4) d15NPN was affected by isotope fractionation during the uptake process, as suggested by negative correlation of d15NPN to 15 NHþ 4 concentration. Positive correlation of d N to temperature was also observed especially for the microplankton fractions, which might have resulted from seasonal variation of d15N of NHþ 4 in estuarine water. (5) The fraction of allochthonous POC, as calculated by the dynamic end member model using d13Cchl, was higher during the autumnal bloom than during the spring bloom, reflecting seasonal change in the river discharge. Taking account of possible variability of d13Cchl relative to the whole algal cell, the maximal uncertainty of 16e22% was expected for the estimates of fractional abundance of allochthonous POC by our model. (6) The size distribution of allochthonous POC transported by the river water was largely biased toward the silt and clay fractions (i.e. <20 mm). (7) Concentration of allochthonous POC in this estuary was principally controlled by conservative mixing between the river water and seawater, though a moderate influence by salt-induced coagulation and/or river flow rate was also suggested. Acknowledgments We are grateful to Drs Yoji Nakajima and Naohiko Ohkouchi (Institute for Frontier Research on Earth Evolution, Yokosuka) for technical advice on isolation and isotopic analysis of chlorophylls. We also thank Drs Keiri Imai, Hideki Fukuda and Ms Nobue Saotome (Ocean Research Institute) for technical assistance with the microscopic examination of phytoplankton, filtration, and fluorometric analysis of chlorophylls, respectively. We also benefited from valuable comments on an earlier version of our manuscript provided by Dr Tim Jennerjahn (Center for Tropical Marine Ecology, Bremen) and two anonymous reviewers. This study was financially supported by CREST, R&D of Hydrological Modeling and Water Resources System (Japan Science and Technology Agency) and The 21st Century COE Program, Biodiversity and Ecosystem Restoration Research Project (Japan Ministry of Education, Culture, Science and Technology). References Altabet, M.A., Francois, R., 2001. Nitrogen isotope biogeochemistry of the Antarctic Polar Frontal Zone at 170  W. Deep-Sea Research II 48, 4247e4273.

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