Sinking dynamics of particulate matter in the subarctic and subtropical regions of the western North Pacific

Author’s Accepted Manuscript Sinking dynamics of particulate matter in the subarctic and subtropical regions of the western North Pacific Chiho Sukigara, Yoshihisa Mino, Hajime Kawakami, Makio C. Honda, Tetsuichi Fujiki, Kazuhiko Matsumoto, Masahide Wakita, Toshiro Saino

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S0967-0637(18)30158-4 https://doi.org/10.1016/j.dsr.2018.11.004 DSRI2977

To appear in: Deep-Sea Research Part I Received date: 15 May 2018 Revised date: 31 October 2018 Accepted date: 16 November 2018 Cite this article as: Chiho Sukigara, Yoshihisa Mino, Hajime Kawakami, Makio C. Honda, Tetsuichi Fujiki, Kazuhiko Matsumoto, Masahide Wakita and Toshiro Saino, Sinking dynamics of particulate matter in the subarctic and subtropical regions of the western North Pacific, Deep-Sea Research Part I, https://doi.org/10.1016/j.dsr.2018.11.004 This is a PDF file of an unedited manuscript that has been accepted for publication. As a service to our customers we are providing this early version of the manuscript. The manuscript will undergo copyediting, typesetting, and review of the resulting galley proof before it is published in its final citable form. Please note that during the production process errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.

Sinking dynamics of particulate matter in the subarctic and subtropical regions of the western North Pacific

Chiho Sukigara1*, Yoshihisa Mino2, Hajime Kawakami3, Makio C. Honda3, Tetsuichi Fujiki3, Kazuhiko Matsumoto3, Masahide Wakita3, Toshiro Saino

1

Tokyo University of Marine Science and Technology, Tokyo, Japan

2

Institute for Space-Earth Environmental Research, Nagoya University, Nagoya, Japan

3

Japan Agency for Marine-Earth Science and Technology, Yokosuka, Japan

*Corresponding author: Chiho Sukigara, Tokyo University of Science and Technology, 4-5-7,

Konan,

Minato-ku,

Tokyo,

108-8477.

Tel:

+81-3-5463-0595;

Fax:

+81-3-5463-0715. e-mail: [email protected]

Abstract Sinking particles collected by drifting sediment traps at 100 and 200 m depths at two observation sites, K2 and S1, located in the subarctic and subtropical gyres in the western North Pacific, respectively, were fractionated in 5 ranges of sinking velocities

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between 5 and 1000 m d-1 using an elutriation system. The velocity distributions were divided into two distinct types over both sites: type-S containing more particles at the relatively slow end of the velocity range and type-F including a peak in the middle range (50 to 150 m d-1). These distributions showed little change between 100 to 200 m, although fluxes of particulate organic carbon (POC) decreased vertically. The averaged sinking velocities (wpoc) calculated from the velocity distributions of POC were 31±16 and 63±26 m d-1 at K2 and S1, respectively. For S1 particles, a positive correlation was found between wpoc and content of CaCO3. This result indicates that particles containing large amounts of denser CaCO3 sink faster than those containing large amounts of organic matter with low densities. Particles at K2, which were mainly composed of opal and organic matter, did not exhibit a clear relationship between wpoc and the content of denser opal. Instead of opal, wpoc had a positive correlation with 15N of the sinking particles and was small (large) when the surface layer was stratified (well-mixed). The attenuation of POC flux with depth resulted from physical fragmentation of particles by turbulence at K2 and from biological decomposition and fragmentation at S1. Keywords: Particulate sinking velocity, POC flux attenuation, mineral ballast effect, stable nitrogen isotope ratio

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1 Introduction The sinking of particulate matter in the ocean is one of the most important processes by which carbon and other biophilic elements are transported from the surface to the ocean interior, referred to as the “biological pump” (Volk and Hoffert 1985). CO2 is converted to particulate organic carbon (POC) via phytoplankton photosynthesis in the euphotic zone, after which some of the POC sinks and is oxidized to CO2 in the deep ocean (Broecker and Peng 1982). This transport acts to maintain a vertical gradient in the concentration of dissolved inorganic carbon in the ocean and eventually controls the air-sea balance of CO2. An understanding of the factors regulating the biological pump is required to predict the atmospheric CO2 concentration in the future. The flux of sinking particles beneath the euphotic zone decreases with the depth. Martin et al. (1987) suggested that the attenuation of the POC flux could be expressed as a simple power law as a function of depth and the flux at a 100 m depth: 𝑧 −𝑏

𝐹(𝑧) = 𝐹(𝑧0 ) × (𝑧 ) 0

z0 = 100 m

(1)

where z is the depth, F(z) and F(z0) are POC fluxes at depths of z and z0, respectively, and the exponent b is a dimensionless scaling factor that indicates the magnitude of flux attenuation with depth. In general, most of the POC is remineralized while sinking through the upper few hundred meters of the water column (Martin et al. 1987, Berelson

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2001). The exponent b varies from 0.5 (in the Northwest Pacific subarctic gyre) to 1.6 (in the North Atlantic subtropical gyre) when fitting Eq. 1 to the flux data (Lamborg et al. 2008, Marsay et al. 2015). Consequently, variations of the exponent b as well as the primary productivity (i.e. particle formation) result in large spatial variations of the annual POC fluxes into the ocean interior by two to three orders of magnitude (Honjo et al. 2008). Sinking velocity is the crucial parameter regulating the attenuation rate of POC flux with depth (i.e. exponent b), because it determines the retention time of particles within a given layer of the water column. If a sinking particle undergoes bacterial decomposition with a constant rate, more POC would be lost from slower sinking particles due to remineralization occurring for a longer time in the water column. Thus, the sinking velocity of the particles has a large influence on the efficiency of the biological pump (Goutx et al. 2007, Trull et al. 2008). In general, the sinking velocity (w, m s-1) can be expressed using the equivalent sphere diameter (D, m), the particle density (1, kg m-3), the seawater density (2, kg m-3), gravitational acceleration (g, 9.81 m s-1), and the drag coefficient (CD, non-dimensional): 𝑤 = √[

4×(𝜌1 −𝜌2 )×𝐷×𝑔 3𝜌2 ×𝐶𝐷

]

(2)

Assuming a particle Reynolds number of < 2, the sinking velocity is described to the

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following form known as Stokes’ law, 𝑤=

𝑔×𝐷 2 ×(𝜌1 −𝜌2 ) 18×𝜇

,

(3)

where  (g m-1 s-1) is the dynamic viscosity. These equations (Eq. 2 and 3) suggest that larger and denser particles sink faster. However, the application of these formulas is not realistic for estimating the sinking velocity because of the difficulty of measuring the densities of porous particles or particle aggregates as well as uncertainties regarding the assumptions of spherical shape and Reynolds number of actual particles (McDonnell and Buesseler 2010, Miklasz and Denny 2010, Iversen et al. 2010). Most sinking particles are fragile, and their physical characteristics are easily changed during sampling and subsequent treatment of the particles. Several methods have been introduced to assess the in situ sinking velocity of oceanic particles, including the use of particle settling columns (Silver and Alldredge 1981), the net-jet flow system (Ploug and Jørgensen 1999), introducing radioactive material in the sinking particles (Villa-Alfageme et al, 2014), combinations of polyacrylamide gel traps and a video plankton recorder (McDonnell and Buesseler, 2012), the elutriation system (Peterson et al., 2005), a sophisticated sediment trap with indented rotating spheres (Peterson et al., 2005, Armstrong et al., 2009, Alonso-González et al., 2010, Riley et al., 2012), as well as digital imaging systems (Stemmann et al., 2004, Ploug et al., 2010,). These methods

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have revealed that sinking velocities of particles varied from a few m d-1 to 1000 m d-1 (e.g. Peterson et al., 2005). As one of the factors governing the sinking velocities through the ballast effect, mineral components of the sinking particles can play an important role (Armstrong et al., 2001, Ploug et al., 2008, Engel et al., 2009, Lee et al., 2009, Honda and Watanabe, 2010). However, there are not enough data related to the particulate sinking velocity to evaluate its general features in relation to particle dynamics in the ocean. By using the elutriation method herein, we assessed the sinking velocity of particles collected at two contrasting sites in the subarctic (K2: 47oN, 160oE) and subtropical (S1: 30oN, 145oE) gyres, as a part of the “K2S1project” intended to further understand the carbon cycles in both regions with a comparative study of the ecosystem and its biogeochemistry (Honda et al. 2017). Analysis of the chemical properties of the particles as well as the sinking velocities allowed us to examine the effect of mineral ballasts on the POC flux. In addition, we also investigated the decomposition processes of sinking particles by the comparing the biogeochemical and stable isotope data from sediment traps at different depths.

2 Materials and methods

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2.1 Drifting sediment trap experiment A free drifting sediment trap array with particle collectors were deployed during five cruises aboard the R.V. Mirai (MR10-06: Oct.– Nov. 2010, MR11-02: Feb.– Mar. 2011, MR11-03: Apr. 2011, MR11-05: Jul. 2011, MR12-02: Jun.–Jul. 2012) at the subarctic station K2 and the subtropical station S1 (Table 1). Details about the trap deployment and sample collection have been reported separately by Honda et al. (2015). In brief, cylindrical particle traps (Knauer type, Knauer et al. 1979) were mounted to a PVC cross on a drifter at depths of 60, 100, 150 and 200 m as a basic design. To collect enough particles required for the elutriation experiments, additional traps were mounted at 90 and 95 m, and at 180 and 190 m, and samples from these two sets of depths were mixed and used to represent particles at depths of approximately 100 and 200 m, respectively. During cruise MR12-02, we also collected particles from 480 and 490 m, correspond to particles at a 500 m depth. The traps at 200 m at S1 during cruise MR10-06 could not be recovered. Before elutriation, a portion of the sinking particles was used for morphological observation (sizes and shapes) using a binocular stereomicroscope. The exponential b (Eq. 1) values of the POC used in this study are those reported by Honda et al. (2015). In this study, we used the chemical compositions (concentration of opal, CaCO3, and lithogenic materials) of sinking particles reported by

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Honda et al. (2015). Concentrations of Si, Ca, and Al were measured using ICP-AES (PerkinElmer Optima 3300DV), and the fraction of each component was estimated based on crustal ratios. 2.2 Elutriation The particulate samples trapped at 100 and 200 m (and 500 m during MR12-02) were fractionated by their sinking velocities using the elutriation method (Peterson et al., 2005). Our system (Fig. 1) consisted of two polycarbonate columns with 400 mm (length) × 45 mm (diameter) for the narrow column and 300 mm (length) × 85 mm (diameter) for the wide column. For the both columns, an upward flow of filtered or artificial seawater at around 20 ℃ was maintained by a peristaltic pump with a constant flow rate of 57 mL min-1. Trapped particles at each depth were introduced via a funnel into the narrow column, shown on the left in Fig. 1. The particles that were retained at the bottom of the column sank at velocities above the threshold determined by the water flow rate and cross-section of the narrow column. These particles were classed as the ‘Fast (sinking) Fraction’. Slower sinking particles were subsequently swept into the wide column and were fractionated again by their velocities and the threshold related to the flow rate and column cross-section. The retained particles in this column were classified as the ‘Middle Fraction (IV)’, while those swept into the tank were classified

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as the ‘Slow Fraction (V)’. The ‘Fast Fraction’ particles were retrieved and further divided into three fractions (I+II+III) using the same columns but with a faster flow rate (675 mL min-1). Finally, we divided trapped particles into five sinking velocity fractions, as shown in Table 2. The median of fraction I was set to be 750 m d-1, which is an intermediate value between the minimum velocity for the narrow column (500 m d-1) and the maximum (about 1000 m d-1) reported in the literature (Peterson et al., 2005, Lee and Niiler, 2005, Trull et al., 2008). The particle fractionation by elutriation was conducted on board for the samples collected during MR10-06 (K2 and S1) and MR11-02 (S1), while fractionation was performed onshore for the other cruises’ samples, which were transported to a laboratory after poisoning with 10% formalin. Before commencement of both the shipboard and laboratory elutriation, zooplankton swimmers were removed under a binocular stereomicroscope. To test the repeatability of our elutriation, we split the particles collected at 100 m during cruise MR11-03 and at 200 m during MR11-05 at K2 into three and five subsamples, respectively, conducted the velocity fractionation for each and compared the POC distributions with the velocity fractions (Fig. 2) among them. Triplicate measurements with samples from MR11-03 showed a relatively large variability of the slower fractions III, IV, and V (Fig. 2a). However, less variability was observed in

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quintuplicate measurements with the samples from MR11-05. Such differenced in the repeatability can be attributed to the distinct sample properties, in which fragile particle aggregations were fragmented easily due to the upward water flows during elutriation. The variabilities in these replicate experiments were considered to be measurement errors when the averaged sinking velocity of the particles (wPOC) (see details of the calculation explained in Section 3.3.1) was estimated. Estimates of the wPOC errors were ± 9.2 m d-1 and ± 6.8 m d-1 for particles collected during cruise MR11-03 and MR11-05, respectively. We believe that the elutriation method provides important information about the sinking velocities of particles. However, it is obvious that this system cannot completely simulate sinking particles in the ocean. First, particles must aggregate at the bottom of the sampling cup during the trap experiment and subsequently must also change in size and shape during the sample transport and treatment processes. Furthermore, in the elutriation experiments, the particles would be fragmented by the forceful flow when the particles are introduced into the elutriation columns with seawater and the direction of flow changes from downward to upward at the bottom of the column. Consequently, the sinking velocities of particles measured by this method tend to be underestimated because of the particle fragmentation process.

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2.3 Analyses of particulate organic carbon and nitrogen, and isotopes Particles in each velocity fraction were collected on pre-weighed, pre-combusted glass-fiber filters (0.7 m pore size, 25 mm diameter, GF/F, Whatman®). All filters were kept in a freezer at -20 ℃ until analysis. The particles on the filters were dried in a desiccator, weighed, and exposed to HCl fumes for 12 h to remove inorganic carbonates prior to measuring POC. The contents of organic carbon (Corg%) and nitrogen (N%) and the nitrogen isotopic composition (δ15N) of the particulate samples were measured using a continuous flow isotope ratio mass spectrometer (Delta PLUS®, Thermo Fisher Scientific) fitted with an elemental analyzer (NC-2000®, CE Instruments) interfaced with a ConFlo II® (Thermo Fisher Scientific). The analytical precision based on the replicate analyses of δ15N was ±0.2‰ for our measurements. 2.4 Biological oxygen consumption rates To assess the in situ decomposition of organic matter in the water column at K2 and S1, we estimated biological oxygen consumption rates (OCR) from dark bottle incubations conducted during cruises MR11-02 and MR11-05. OCR was measured as changes in the dissolved oxygen (DO) concentration before and after incubation. Water samples at 100 and 200 m were collected in WOCE-type DO bottles (~ 100 mL) and incubated at stable temperatures (4 and 20 ℃ for samples at K2 and S1, respectively) in the dark. During

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the 1-week incubation experiments, DO measurements were conducted on Day 0, Day 7, and on 2 different days in between, in which triplicate bottles were treated with Winkler reagents. DO concentrations were determined by a high precision titration unit (MPT Titrino 798; Metrohm Shibata) according to the Winkler method (Japan Meteorological Agency, 1999). OCR was calculated from the slope of the regression of the variation of DO concentration with incubation time.

3 Results and discussion 3.1 Vertical changes of POC fluxes and contents POC fluxes in the subarctic K2 at 100 m and 200 m varied seasonally from 15.1 to 50.0, and from 13.4 to 44.4 mg C m-2 d-1, respectively (Fig. 3a). Decreasing rates from 100 to 200 m ranged from 40 to 50% (of the flux at 100 m) for most cruises, with low rates observed during MR10-06 (from 15.1 to 13.4 mg C m-2 d-1) and MR11-05 (from 42.9 to 44.4 mg C m-2 d-1). POC flux at 500 m during MR12-02 decreased by 30% relative to the flux at 200 m. The POC fluxes in the subtropical S1 at 100 and 200 m varied from 13.8 to 59.2, and 16.1 to 36.9 mg C m-2 d-1, respectively (Fig. 3b). The decreasing rates from 100 to 200 m ranged from 23% to 57%, and was 50% from 200 to 500 m during MR12-02. Kobari et al. (2016), who investigated carbon flux by mesozooplankton

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communities in the same regions (K2 and S1), estimated that the contributions of fecal pellets of mesozooplankton to the entire carbon flux varied from 5 to 35% and 1 to 4% at K2 and S1, respectively. Carbon fluxes due to fecal pellets during MR11-02 and MR11-03 at K2 and MR11-02 at S1 were remarkably reduced from 100 to 200 m. These decreases corresponded to less than 30% of the loss of the entire POC fluxes in the same layer. It is suspected that these flux attenuations with depth were attributed to bacterial decomposition of the organic matter of the particles while sinking. To test this hypothesis, the Corg% of sinking particles was compared between 100 and 200 m, although available data was limited. In addition, the N% of the sinking particles showed a similar trend to that of Corg%. The Corg% at K2 ranged 13 to 20% at 100 m and was around 19% at 200 m (Fig. 3c). Vertical changes between both depths were less than ±5%. In the 500-m sample from MR12-02, the Corg% decreased by 3% relative to that at 200 m. These small changes with depth do not suggest a predominant contribution of bacteria to the POC flux attenuation at K2. Meanwhile, the Corg% at 100 and 200 m at S1 varied from 13 to 28% and from 16 to 22%, respectively (Fig. 3d). A relatively large decrease (7%) from 100 to 200 m was found during MR11-05 when the POC flux attenuated significantly (by 57%), suggesting that the removal of organic matter via

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bacterial decomposition was predominant. During MR12-02, however, the Corg% did not decrease clearly between 200 and 500 m, even with the large flux attenuation of 50%. The estimated OCR at K2 during two cruises (MR11-02 and MR11-05) ranged from 0.12 to 0.14 μmol O2 kg-1 d-1 at 100 m and 0.01 to 0.03 μmol O2 kg-1 d-1 at 200 m (Table 3), likely indicating less seasonal variation but significant changes with depth. The lower OCRs estimated at 200 m were within the margin of analytical error. At S1, the OCRs were much larger than those at K2, ranging from 0.43 to 0.51 μmol O2 kg-1 d-1 at 100 m and 0.27 to 0.51 μmol O2 kg-1 d-1 at 200 m. The summer OCR at 200 m (0.27 μmol O2 kg-1 d-1) was comparable to the estimate of 0.23 μmol O2 kg-1 d-1 for the subtropical Mode Water between March and July, reported by Sukigara et al. (2011), based on time-series DO variations obtained from a float equipped with an oxygen sensor. The lower heterotrophic prokaryotic production and prokaryotic respiration at K2 than at S1 (Uchimiya et al., 2015) also suggest that there was less biological decomposition of organic matter at the former. It should be noted that our OCR values account for dissolved organic carbon and fine particles respiration and cannot be directly compared to those for sinking particles. However, considering that the bacterial production rate depends on the ambient temperature (Uchiyama et al., 2015) and the

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decomposition of sinking particles occurs due to particle-attached bacteria, the OCR values are useful for discussing the attenuation of POC fluxes. Oxygen consumption rates estimated from the carbon demands of mesozooplankton (Kobari et al., 2016) corresponded to about 15% of the OCR at K2 (and a few percent at S1). These values are not negligible but would not significantly change the OCR trends. 3.2 Velocity distributions of POC in the sinking particles measured by the elutriation method At K2 during four of five cruises, more POC in the sinking particles was classified as the relatively slow fractions III, IV, and V (i.e. <150 m d-1) than as the faster fractions I and II (Fig. 4b-e). We define such distributions as “type-S (Slow)”. However, during MR10-06, the POC peak was observed in fraction III (50-150 m d-1), defined as “type-F (Fast)”, and the POC contributions of fractions I and II were relative high compared with those observed during the other cruises (Fig. 4a). At S1, type-F was observed during MR11-02 and MR12-02 (Fig. 4g and j), while type-S was found during MR10-06, MR11-03, and MR11-05 (Fig. 4f, h, and i). During each cruise in both K2 and S1, similar (or nearly the same) distributions were observed at 100 and 200 m. At K2, the Corg% for each velocity fraction mostly showed the same values for the particles collected at 100 and 200 m (Fig. 5a-d). However, a significant increase with

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depth was seen in some fractions (I and II during MR11-02 and all but III during MR12-02), the reason for which is unknown. In contrast, for particles at S1, an obvious Corg% decrease with depth was seen in both the faster and slower fractions (Fig. 5e-g), especially during MR11-05 when the largest bulk POC attenuation was found (Fig. 3b). Note that the apparently larger Corg% decreases in the faster fractions I and II did not account for bulk Corg% decrease of 7% with depth during MR11-05 (Fig. 4b) due to their lower POC contributions (<0.1, Fig. 4i). Alternatively, the Corg% in the slower fractions III-V decreased. During MR12-02, the fraction III particles showed relatively low Corg% (<15%), even for the largest POC contribution (i.e. type-F, Fig. 4j), and such Corg% was nearly constant against depth. These results at K2 and S1 revealed that two patterns of velocity distributions, type-S and type-F appeared over both regions and that these distributions were almost unchanged between 100 and 200 m even with significant losses of POC flux with depth. Given a constant rate of POC decomposition among all the velocity fractions, the slower particles (the fraction V) would lose more POC than the faster particles during their sinking through the same water column. However, our results of Corg% in the velocity-fractioned particles contradicted the preferential POC loss in a specific fraction. They suggested that variable-rate POC loss would occur among the particles that sink

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with different velocities. Alternatively, other processes independent of the particle sinking velocity would correspond to the POC attenuations with depth. 3.3 Factors controlling sinking velocity of particles 3.3.1 Estimation of an average sinking velocity To address the causes of different velocities between sites and seasons, we first estimated the average sinking velocities (w) of bulk particles and compared them with other particulate properties. Logarithmic averages of the sinking velocities (wPOC) were calculated based on the relative POC fluxes in five velocity ranges (fi) and their medians (wi) (Table 2) as follows: 𝑤𝑃𝑂𝐶 = exp[∑5𝑖=1{𝑓𝑖 × ln(𝑤𝑖 )}]

(4).

At K2, the average sinking velocity of POC (wPOC) estimated from velocity distributions of POC (Fig. 4) varied from 28 (100 and 500 m at MR12-02) to 110 m d-1 (200 m at MR10-06) (Table 4). As noted above, the fluid (seawater) temperature in this elutriation was at about 20 ℃, which was quite higher than the in situ temperature at K2 of 1.5 and 3.5 ℃ at 100 and 200 m, respectively (Table 1). The viscosity of seawater decreases as temperature increases. The relative ratios of the viscosity at 20 ℃ to those at 1.5 and 3.5 ℃ are 0.60 and 0.64, respectively (Sharqawy et al., 2010). We used these values in Eq.3 to correct wPOC at K2 for the temperature effect on the viscosity (Table 4). The

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corrected wPOC (i.e. the velocity in situ temperature) at K2 varied from 17 (MR12-02, 100 m) to 70 m d-1 (MR10-06, 200 m) with an average of 31 ± 16 m d-1 (mean ± SD). The estimated wPOC during MR10-06 with the type-F distribution were over 50 m d-1, whereas it ranged below the overall mean during other cruises with the type-S distribution. Comparing 100 and 200 m depths for each cruise, wPOC increased significantly with depth during MR10-06, but vertical differences for other cruises were within the measurement error (± 9.2 m d-1). At S1, wPOC varied from 34 m d-1 (MR11-05, 100 m) to 101 m d-1 (MR11-02, 200 m), (Table 4) with an average 63 ± 26 m d-1, which was significantly larger than that at K2 (t-test, p < 0.005). The derived wPOC values during MR11-02 and MR12-02 with type-F distributions were over 80 m d-1, while wPOC ranged 30 to 40 m d-1 during MR11-03 and MR11-05 with the type-S distributions. Differences in wPOC between 100 and 200 m for each cruise were small. The estimated wPOC at 500 m during MR12-02 was smaller (70 m d-1) than that at 200 m (85 m d-1). Honda et al. (2013) reported the sinking velocities of trapped particles at 500 m as 22 to 46 m d-1 at K2 and 26 to 71 m d-1 at S1, which were inferred from the time lag until the first detection of the radiocesium signal emitted from the Fukushima Daiichi Nuclear Power Plant accident. Our average wPOC of 31 m d-1 (at 5 ℃) at K2 and 63 m d-1 (at

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20 ℃) S1 (Table 4) are within the velocity ranges reported by Honda et al. (2013), which provides additional support for the reliability of our measurements. We conclude that our method gave appropriate values of particle sinking velocities in the study areas. 3.3.2 Comparisons between chemical or biological parameters and wPOC The chemical composition of particles affects their sinking velocities by controlling their densities. Dense inorganic ballasting materials, such as CaCO3, opal, and lithogenic materials (LM), is expected to increase particle-sinking velocities. Klaas and Archer (2002) reported that organic carbon in the deep sea was carried by dense CaCO3, while Honda and Watanabe (2010) claimed the importance of biogenic opal as a ballast for particulate organic matter transport in the Western Pacific Subarctic Gyre. Therefore, we compared our estimated wPOC (at 20 ℃) at both sites with the relative compositions of CaCO3, opal, and organic matter (OM) in the bulk particles reported by Honda et al. (2015). Note that CaCO3 contents were generally higher at S1 than K2, but the opal contents were higher at K2 and the OM contents had similar ranges (30 to 60 %) at both sites. As for LM, relatively low contents (<5 %) were found at both sites during all the cruises except for MR11-03, whose values exceeded 20 % at S1. At K2, the two largest wPOC (87 and 110 m d-1 at 20 ℃ during MR10-06, Table 4) coincided with higher CaCO3 contents (40 %, Fig. 6a). Microscopic observations also revealed that relatively large

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amounts of foraminiferal shells were included in fraction I during MR10-06. These CaCO3 materials would contribute to a larger wPOC. During MR11-05 and MR12-02, diatoms were predominant (25 to 30% of Chl a, Fujiki et al. 2014). However, the wPOC values were relatively small (~34 m d-1) despite higher opal contents (40 to 60 %) (Fig. 6b). Particles at S1 were composed mainly of OM and CaCO3 (Fig. 6a, c), in which their sum accounted for ca 85-90% of the total mass on average (Honda et al. 2015). The wPOC values showed a positive correlation with the CaCO3 content (y = 222.3x − 20.1, r = 0.93, p < 0.01, dashed line in Fig.6a) and a negative correlation with OM (y = -261.4x − 185.8, r = 0.79, p < 0.01, in Fig.6c). As noted above, the large wPOC values observed during MR11-02 and MR12-02 were attributed to the type-F velocity distributions with a peak POC contribution of fraction III (Fig. 4g and j). During MR12-02, the Corg% in fraction III was < 15% and lower than those (26-32%) found during other cruises (Fig. 5e-g). This low Corg% implies that the fraction III particles during MR12-02 contained a large amount of CaCO3, and, conversely, these particles would sink potentially with velocities in the range of 50-150 m d-1 due to the effect of denser CaCO3. Based on the results at K2 and S1, the CaCO3 content positively affected the wPOC. However, this effect is obscured for particles with CaCO3 contents <20 %, such as

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particles collected at K2 (Fig. 6a). Meanwhile, we found no such effect for opal in Fig. 6b, where the wPOC values were relatively constant or decreased slightly with opal contents increasing from 20 to 60 %, even though these changes of opal content would control the bulk particle density. This implied that other factors might counteract the effect of opal ballast at K2. We discuss the factors that affect the sinking velocities at K2 in detail below. Stokes’ law (Eq. 3) predicts that small particles tend to sink slowly because of their large surface areas relative to their volumes, and vice versa. Unfortunately, there are no data available on the sizes of trapped particles. Alternatively, we used the size distributions of phytoplankton during each cruise (Fujiki et al., 2014, 2015) to compare with the wPOC values (Figs. 7a-c). During most of cruises, the small plankton (<3 μm) were dominant at both sites, with more contributions at the subtropical S1. However, large plankton (>10 μm) contributed more at the subarctic K2. Compared to the size distributions at both sites, the wPOC values did not show any significant correlation. This implied that the primary producer’s size would not determine the size of sinking particles, and that the size distributions of particles would have less influence on the velocity. 3.3.3 Relationship between wPOC and 15N of sinking particles

21

If the variation of sinking velocity was not due to the chemical composition, the size and/or shape of sinking particles might determine their velocities. Unfortunately, we did not observe these under a microscope after velocity fractionation. However, the sizes and shapes may vary due to decomposition and transformation of POC. We measured 13Corg (not shown) and 15N of trapped particles because these values vary due to isotope fractionation when organic matter is involved in biological and chemical reactions. We examined how the wPOC values were related to the 15N values of the particles because 15N varies depending on the nitrogen sources and their availabilities when the particles formed (Wada and Hattori 1991, Altabet and François 1994, Mino et al. 2002 etc.). 15N of the particles trapped at both sites ranged from 1.2 to 7.1‰, with larger seasonal variations at K2 (Fig. 8). Interestingly, a positive correlation was found between wPOC and particulate 15N at K2 except in the data collected during MR10-06 with the two largest wPOC values (y = 0.35x – 3.62, r = 0.80, p < 0.05, dashed line in Fig. 8b). Among these data, relatively high 15N (> 6‰) and large wPOC values (42-51 m d-1) were observed in winter/spring (MR11-02 and MR11-03) when surface mixing layers reached depths greater than 100 m (Table 1). On the other hand, lower 15N (<4‰) and smaller wPOC (~34 m d-1) were found in the summer (MR11-05 and MR12-02) when the

22

mixing layers were shallower (27-39 m, Table 1) with relatively low nitrate concentrations. Mino et al. (2016) proposed that such seasonal 15N variations cannot be explained by isotopic fractionation during phytoplankton nitrate consumption and attributed it alternatively to the light-controlled ammonium (NH4+) consumption by nitrifiers. In this process, preferential 14NH4+ oxidation by nitrifiers in the deeply mixed water would cause higher 15N values of the NH4+ pool associated with extremely large isotopic fractionation (-38‰, Casciotti et al. 2003). This enriched

15

N signal was

imparted to phytoplankton by uptake of heavy NH4+, even with isotopic fractionation (-27‰, Wasser et al. 1998), which formed particulate nitrogen with higher 15N in winter. Such NH4+ oxidation was suppressed in the bright, stratified water, which resulted in relative low particulate 15N in summer. Based on their hypothesis, the 15N signature at K2 would reflect the seasonal change in water column mixing. Therefore, our derived wPOC vs 15N correlation at K2 implies that water column mixing might affect the particle-sinking velocity. At K2, where diatoms are dominant (Table 4), particle aggregation occurs in the upper column because of transparent exopolymer particles (TEPs) (Alldredge et al, 1993, Passow et al., 2001, Engel et al., 2004, Mari et al., 2017). Mari et al. (2017) suggested that TEPs with low densities contribute to the buoyancy of POC and decrease sinking

23

velocity. Loose and fluffy aggregates might sink slowly due to their low densities and/or large drag (Bach et al., 2016). In this scenario, the aggregation that lowers the sinking velocities of bulk particles would be enhanced in stable, stratified water, while it would be restricted by active turbulent mixing, which would fragment the aggregates. Mino et al. (2016) reported higher Corg:N mole ratio in the sinking particles compared with the suspended ones at K2 throughout a year, and they attributed this to particle aggregation via TEPs. Moreover, this Corg:N ratio was elevated to ~8.8 in the summer from 7.2 in the winter, supporting our speculation that enhanced particle aggregation occurs under stratified conditions. Thus, it is likely that the effect of water mixing on the formation of slowly sinking aggregates partially determines the seasonal velocity change for the bulk particles trapped at K2, which results in slower and faster wPOC in the summer and winter/spring, respectively. If these mechanisms are robust, the positive relationship between wPOC vs 15N derived herein allow estimates to be made for sinking velocities of particles at K2 from the 15N analysis. Conversely a negative correlation was found at S1 between wPOC and δ15N (y = – 0.028x + 5.36, r = 0.68, p < 0.05, Fig. 8a). At the oligotrophic S1 region, trapped particulate δ15N seasonally varied depending strongly on nitrate availability in the euphotic zone, i.e. it was governed by the isotopic fractionation associated with algal nitrate

24

consumption (Mino, personal communication). Lower δ15N signals (< 4‰) found during MR11-02 and MR12-02 reflect the relatively high availability of nitrate, which was supported by the larger POC fluxes measured during these cruises compared to those measured during the others (Fig. 3b). Enhanced growth of coccolithophores under such conditions likely led to higher CaCO3 contents (> 0.4) and thereby larger velocities (> 80 m d-1) of the trapped particles (Fig. 6a). However, with less nitrate availability during the other cruises inferred by the higher δ15N values (>4‰, Fig. 8a), non-calcifying picoeukaryotes as well as prokaryotic Prochlorococcus were dominant (Fujiki et al., 2015), resulting in moderate sinking velocities (< 60 m d-1) of the particles. 3.4 Relationship between wPOC and POC flux attenuation coefficient with depth (exponent b) To address the attenuation of POC fluxes, we examined the relationship between wPOC and the exponent b, the scaling factor for POC flux attenuation with depth. We used the exponent b reported in Honda et al. (2015), in which they applied the power law function (Martin et al. 1987) to POC fluxes at 60-200 m at K2 and S1 during each cruise. The sinking velocities, wPOC, corrected for in situ temperature were used. Generally, if the total POC decay rate (k, d-1) is constant regardless of processes, a

25

negative relationship is expected between wPOC and the exponent b (Eq.5; Sarmiento and Gruber, 2006). 𝐹(𝑧) = 𝐹(𝑧0 ) × exp [𝑤

𝑘

𝑃𝑂𝐶

× (𝑧 − 𝑧0 )] (z0 = 100, z = 200)

(5)

As seen in Fig. 9, however, no such a trend was found in the data from either or both site, which implied that the decay rate k varied seasonally as well as regionally. At K2, the exponential b varied largely from 0.2 to 1.0, while wPOC had a small range when excluding the data obtained during MR10-06. This implied that there was more than a 20-fold change in k between the maximum (0.21 d-1) and the minimum (0.01 d-1) during MR11-02 and MR11-05, respectively. The question remains: what causes such variation in decay rate of POC? The vertically constant Corg% of the bulk particles indicate insignificant biological OM decomposition at K2 (Fig. 4a). In that case, particle fragmentation by an abiotic, physical process could cause the POC flux attenuation, in which fragmented smaller particles (or aggregates) would remain suspended. During MR11-02 in the winter, as noted in Sect. 3.3.3, active turbulence would act to restrict the formation of slow-sinking, fluffy aggregates (which results in larger wPOC), but this physical forcing contributed to a larger decay rate k by altering the state of the formed aggregates from sinking to suspended. The smaller k during the summer (MR11-05) implied such physical fragmentation rarely occurred through the stable water column

26

and consequently, POC would be transported relatively slowly (34 d-1), with less decay with depth. Limited heterotrophic OM decomposition during summer with the diatom dominance was possibly attributed to opal protection from decomposition as well as low temperature (Bach et al., 2012, Iversen and Ploug, 2013). Relatively large and uniform b values (0.9 – 1.2) were found at S1, even though there was a three-fold variation in wPOC, indicating larger k values for the particles that sank faster and vice versa. The largest two k values of 0.7 and 0.5 d-1 were found during MR11-02 and MR12-02 when larger POC fluxes occurred. Because of the nitrate supplies from below inferred by the low δ15N of the trapped particles, the resulting active mixing likely enhanced particle fragmentation, which would cause larger k values, as was found for K2, in addition to elevated bacterial activity due to increased substrate availabilities (Uchimiya et al., 2015).

4. Conclusion Our investigations of the sinking velocity distributions of particles by elutriation revealed that it varied regionally (subarctic K2 vs subtropical S1) and seasonally. These variations were attributable to changes in not only the chemical composition of the particles (i.e. CaCO3 contents) but also the ambient water stability, which affects the

27

morphology of the particles (aggregates vs fine particles). As a result, there were empirical relationships between the average velocity wPOC with CaCO3 contents at S1 and with δ15N for particles at K2, which enabled us to evaluate the factors controlling wPOC from the elemental and isotopic analyses of particles. By applying these empirical relationships to particles collected by time-series sediment traps in this region (e.g. Honda et al., 2018), CaCO3 contents and δ15N could potentially be used to reconstruct seasonal changes in upper ocean stratification. We also addressed variations in the POC decay rate (k) from wPOC and flux attenuation coefficients (exponent b) at both sites, which suggest that POC could be removed from sinking particles by physical fragmentation (caused by turbulence) as well as biological decomposition. In conclusion, Table 5 shows average properties of POC export to the mesopelagic zone at both sites, suggesting that POC is exported faster but undergoes faster decay at S1 and vice versa at K2. This difference is ultimately attributed to the accelerated wPOC due to the CaCO3 ballast at S1 and the restricted OM decomposition by opal protection and low temperatures at K2. Meanwhile, it is postulated that seasonal water stability/mixing could influence the growth/fragmentation of aggregates, thereby affecting wPOC and k. Such processes should be incorporated into marine ecosystem models to predict oceanic carbon sequestration.

28

Acknowledgement We thank the captain, crew, and scientists of R/Vs Mirai of the Japan Agency for Marine-Earth Science and Technology (JAMSTEC). We also thank the onboard technical team from Marine Works Japan Ltd. We are grateful to J. I. Goes and H. R. Gomes of Columbia University for their helpful discussions, suggestions, and critical reading of this manuscript. We also deeply thank to reviewers and editors who provided us with many useful comments. References Alldredge, A. L., Passow, U., Logan, L. B. E. (1993). The abundance and significance of a class of large, transparent organic particles in the ocean. Deep-Sea Research I, 40, 1131-1140. Alonso-González, I. J., Arístegui, J., Lee, C., Sanchez-Vidal, A., Calafat, A., Fabrés, J., Sangrá, P., Masqué, P., Hernández-Guerra, A., Benítez-Barrios, V. (2010). Role of slowly settling particles in the ocean carbon cycle. Geophysical Research Letters, 37, L13608, doi:10.1029/2010GL043827. Altabet, M. A., François, R. (1994). The use of nitrogen isotope ratio for reconstruction of past changes in the surface ocean nutrient utilization. In: Zahn, R. et al. (eds),

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Carbon cycling in the Glacial Ocean: constraints on the ocean’s role in global change. Springer, New York, pp. 281-306. Armstrong, R. A., Lee, C., Hedges, J. I., Honjo, S., Wakeham, S. G. (2001). A new mechanistic model of organic carbon fluxes in the ocean: based on the quantitative association of POC with ballast minerals. Deep-Sea Research II, 49, 219-236, doi:10.1016/S0967-0645(01)00101-1. Armstrong, R. A., Peterson, M. L., Lee, C., Wakeham, S. G. (2008). Settling velocity spectra and the ballast ratio hypothesis. Deep-Sea Research II, 56, 1470-1478. doi:10.1016/j.dsr2.2008.11.032. Bach, L. T., Riebesell, U., Sett, S., Febiri, S., Rzepka, P., Schulz, K. G. (2012). An approach for particle sinking velocity measurements in the 3-400 m size range and considerations on the effect of temperature on sinking rates. Marine Biology., 159, 1853-1864. https://doi:10.1007/s00227-012-1945-2 Bach, L. T., Boxhammer, T., Larsen, A., Hildebrandt, N., Schulz, K. G., Riebesell, U. (2016). Influence of plankton community structure on the sinking velocity of marine aggregates. Global Biogeochemical Cycles, 30, 1145-1165. https://doi: 10.1029/2016GB005372

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Research Part A, 26 (1), 97-108. Kobari, T., Nakamura, R., Unno, K., Kitamura, M., Tanabe, K., Nagafuku, H., Niibo, A., Kawakami, H., Matsumoto, K., Honda, M. C. (2016). Seasonal variability in carbon demand and flux by mesozooplankton communities at subarctic and subtropical sites in the western North Pacific Ocean. Journal of Oceanography, 72, 403-418, doi: 10.1007/s10872-015-0348-7. Lamborg, C. H., Busseler, K. O., Valdes, J., Bertrand, C. H., Bidigare, R., Manganini, S., Pike, S., Steinberg, D., Trull, T., Wilson, S. (2008). The flux of bio- and lithogenic material associated with sinking particles in the mesopelagic “twilight zone” of the northwest and North Central Pacific Ocean. Deep Sea Research II, 55(14-15), 1540-1563. Lee, C., Peterson, M. L., Wakeham, S. G., Armstrong, R. A., Cochran, J. K., Miquel, J. C., Fowler, S. W., Hirschberg, D., Beck, A., Xue, J. (2009). Particulate organic matter and ballast fluxes measured using time-series and settling velocity sediment traps in the northwestern Mediterranean Sea. Deep-Sea Research II, 56, 1420-1436, doi:10.1016/j.dsr2.200811.029. Mari, X., Passow, U., Migon, C., Burd, A., Legendre, L. (2017). Transparent exopolymer particles: Effects on carbon cycling in the ocean. Progress in

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Oceanography, 151, 13-37, http://dx.doi.org/10.1016/j.pocean.2016.11.002. Marsay, C. M., Sanders, R. J., Henson, S. A., Pabortava, K., Achterberg, E. P., Lampitt, S. (2015). Attenuation of sinking particulate organic carbon flux through the mesopelagic ocean. Proceedings of the National Academy of Sciences of the United States of America, 112, 1089-1094. www.pnas.org/cgi/doi/10.1073/pnas. 1415311112. Martin, J. H., Knauer, G. G., Karl. D. M., Broenkow, W. W. (1987). VERTEX: carbon cycling in the northeast Pacific. Deep-Sea Research Part A, 34, 267-285. McDonnell, A. M. P., Buesseler, K. O. (2010). Variability in the average sinking velocity of marine particles. Limnology and Oceanography, 55(5), 2085-2096. McDonnell, A. M. P., Buesseler, K. O. (2012). A new method for the estimation of sinking particle fluxes from measurements of the particle size distribution, average sinking velocity, and carbon content. Limnology and Oceanography: Methods, 10, 329-346. Miklasz, K. A., Denny, M. W. (2010). Diatom sinking speeds: Improved predictions and insight from a modified Stokes’ law. Limnology and Oceanography, 55 (6), 2513-2525. https://doi:10.4310/lo.2010.55.6.2513 Mino Y., Saino, T., Suzuki, K., Marañón, E. (2002). Isotopic composition of suspended

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particulate nitrogen (15Nsus) in surface waters of the Atlantic Ocean from 50ºN to 50ºS. Global Biogeochemcal Cycles, 16, 1059. https://doi:10.1029/ 2001GB001635. Mino Y., Sukigara, C., Kawakami, H., Honda, M. C., Matsumoto, K., Wakita, M., Kitamura, M., Fujiki, T., Sasaoka, K., Abe, O., Kaiser, J., Saino, T. (2016). Seasonal variations in the nitrogen isotope composition of settling particles at station K2 in the western subarctic North Pacific. Journal of Oceanography, 72, 819-836. https://doi: 10.1007/s10872-016-0381-1 Passow, U., Shipe, R. F., Murray, A., Pak, D. K., Brzezinski, M. A., Alldredge, A. L. (2001). The origin of transparent exopolymer particles (TEP) and their role in the sedimentation of particulate matter. Continental Shelf Research, 21, 327-346. Peterson, M. L., Wakeham, S. G., Lee, C., Askea, M. A., Miquel, J. C. (2005). Novel techniques for collection of sinking particles in the ocean and determining their settling rates. Limnology and Oceanography: Methods, 3, 520-532. Ploug, H., Jørgensen, B. B. (1999). A net-jet flow system for mass transfer and microsensor studies of sinking aggregates. Marine Ecology Progress Series, 176, 279-290. Ploug, H. Iversen, M., Fischer, G. (2008). Ballast, sinking velocity, and apparent

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40

Captions Table 1 Trap deployments and hydrographic information. K2

S1 Surf Euph

Cruise

Deploy

Durat

T(10

T(20

ID

ment

ion

0m)

0m)

Surf

ace

otic

Deploy

Durat

T(10

T(20

ment

ion

0m)

0m)

mixi

layer

ng

Euph

ace

otic

mixi

layer

ng

layer (UTC)

MR10

o

(days)

( C)

( C)

(m)

(m)

3.0

1.3

3.1

44

49

28-Oct-

-06

10

MR11

25-Feb-

-02

11

MR11

19-Apr11

MR11

01-Jul-1 1

MR12

11-Jun-

(days)

( C)

( C)

(m)

(m)

3.1

18.9

17.6

89

55

3.0

18.2

17.8

65

201

3.1

17.9

16.8

65

49

4.3

18.9

17.6

90

20

3.0

17.6

17.1

74

14

15-Feb1.8

3.4

64

105 11 28-Apr-

1.4

3.7

58

117 11 25-Jul-1

5.0 -05

(UTC)

1.6

3.6

39

27 1 28-Jun-

3.0

1.5

3.6

44

39

12

12

Table 2 Range and median values of sinking velocity in each fraction. I

II

III

IV

V

> 500

150 - 500

50 - 150

15 - 50

5 - 15

750

325

100

32.5

10

Sinking velocity (m d-1) wi (m d-1)

o

10

3.0 -03

layer o

08-Nov-

4.0

-02

o

41

Table 3 Oxygen consumption rates (OCR) at 100 and 200 m at Stations K2 and S1. MR11-02 (winter) -1

-1

MR11-05 (summer)

OCR [mol kg d ]

OCR [mol kg-1 d-1]

100 m

0.14

0.12

200 m

< D.L.*

< D.L.

100 m

0.51

0.43

200 m

0.51

0.27

K2

S1

*D.L.: the detection limit

Table 4 Averaged sinking velocities at Stations K2 and S1. wPOC based on POC wPOC based on POC fraction under 20oC (m d-1) 100m

200m

87

average

fraction under in situ o

average

in 100,

temperature (5 C) (m

in 100,

200m

d-1)

200m

500m

Dominant species

100m

200m

500m

110

52

70

51

49

30

31

diatoms, pelagophytes**

33

42

20

27

cryptophytes**

34

34

20

22

diatoms**

28

34

17

22

56

-

87

101

K2 foraminifera*,

MR10-06

chlorophytes**

(autumn) MR11-02 (winter) MR11-03 (spring) MR11-05 (summer) MR12-02 29

49±25

18

31±16

diatoms**

(summer)

S1 MR10-06

prochorococcus, prymnesiophytes***

(autumn) MR11-02

diatoms, chrysophytes, prymnesiophytes***

(winter)

42

MR11-03

prymnesiophytes, 38

45

chrysophytes***

(spring)

prochorococcus, MR11-05

34

38

chrysophytes, prymnesiophytes***

(summer)

prochorococcus, MR12-02

87

85

70

63±26

chrysophytes, prymnesiophytes***

(summer)

Table 5 Comparison of averaged sinking velocity (wPOC), POC attenuation coefficient (b), and degradation coefficient (k) at K2 and S1. wPOC

b

k

23.6 ± 17.4

0.60 ± 0.40

0.17 ± 0.16

64.3 ± 30.0

1.04 ± 0.11

0.47 ± 0.24

(m d-1) K2 (Subarctic) n=5 S1 (Subtropical) n=4

Fig. 1 Schematic diagram of the elutriation used to measure the velocities of sinking particles. The narrow and wide columns were filled with filtered or artificial seawater and the flow rate was controlled by a peristaltic pump. The sediment trap sample was introduced into this stream of flowing filtered or artificial seawater via a funnel to the left. Material with a sinking velocity faster than that of the upward water flow in the column is collected at the bottom of the narrow column (fast fraction), while those with

43

sinking velocities lower than the upward flow travel to the wide column. At the end of elutriation, the water flow was stopped and particles with sinking velocities faster than that of the upward water flow were removed from the two columns (Fast fraction and Middle fraction) and the particles that were not collected in the two columns were collected in a tank, seen on the right (Slow fraction). This provided us with three fractions of particles with different sinking velocities. Particles collected in the narrow column were further separated into three fractions using a faster flow rate. Thus, we obtained 5 fractions of particles, each with a sinking velocity range and a median listed in Table 1. Fig. 2 Results of the elutriation experiment to test for repeatability. (a) Particles collected at 100 m at K2 during MR11-03 cruise and separated into 3 parts and (b) particles collected at 200 m at K2 during MR11-05 cruise and separated into 5 parts. The total amount of POC in each test was 4.7, 2.7, and 2.4 mg C in MR11-03 and 5.1, 1.7, 1.5, 1.4, and 1.3 mg C in MR11-05. Fig. 3 Seasonal variations of POC fluxes at (a) K2 and (b) S1 and organic carbon content at (c) K2 and (d) S1. Fig. 4 Relative POC flux from elutriation experiments at K2: (a) MR10-06 (autumn), (b) MR11-02 (winter), (c) MR11-03 (spring), (d) MR11-05 (summer), and (e) MR12-02

44

(summer), and at S1: (f) MR10-06 (autumn), (g) MR11-02 (winter), (h) MR11-03 (spring), (i) MR11-05 (summer), and (j) MR12-02 (summer). Fig. 5 Organic carbon content in each fraction at K2: (a) MR11-02 (winter), (b) MR11-03 (spring), (c) MR11-05 (summer), and (d) MR12-02 (summer), and at S1: (e) MR11-03 (spring), (f) MR11-05 (summer), and (g) MR12-02 (summer). Fig. 6 Relationships between sinking velocity and (a) calcium carbonate (CaCO3) content, (b) opal content, and (c) organic matter content of particles collected at 100 and 200 m at Stations K2 and S1, with dashed lines denoting the regression lines. These regression lines were derived between sinking velocities and CaCO3 and organic matter contents for S1. Fig. 7 Relationships between averaged sinking velocity at 100 and 200 m and compositions of (d) large size (> 10 m), (e) medium size (3 to 10 m), and small size (< 3 m) phytoplankton. Fig. 8 (a) Relationship between averaged sinking velocity at 100 and 200 m and nitrogen isotope ratio (15N) of sinking particles and (b) an enlarged view of the hatched area in (1) from 10 to 40 m d-1 on the x-axis. Fig. 9 Relationship between averaged sinking velocity (wpoc) at 100 and 200 m and exponential b at stations K2 and S1. Gray lines indicate wPOC and b variations when the

45

POC decay rate k is 005 and 0.1 to 0.8 d-1. The black circle indicates the MR10-06 data with relatively high wPOC values likely due to foraminiferal shell included in the trapped particles.

46

47

48

49

50

Highlight



Sinking velocity of particulate matter measurement by an elutriation system.



Particles containing large amount of CaCO3 sink faster.



Particles in the subarctic region sink faster in the well-mixed season.

51