²¹⁰Po-²¹⁰Pb Disequilibrium in the Western North Pacific Ocean: Particle Cycling and POC Export

Introduction

As the Earth’s largest carbon reservoir (IPCC in Climate Change, 2013), the marine ecosystem plays a fundamental role in regulating the atmospheric CO₂ concentration and buffering the effects of global climate change (Sabine, 2004) by assimilating carbon from the atmosphere via dissolution and photosynthesis in the upper ocean. Even small changes in the magnitude of the downward transport of carbon in the upper ocean can have serious impacts on oceanic carbon sequestration and atmospheric CO₂ concentration (Kwon et al., 2009) and thus on global climate change (including global warming). It has been predicted that without the marine biological carbon pump (BCP), the atmospheric CO₂ concentration could be approximately 50% higher than its current value (Sanders et al., 2014).

To fully understand how the oceans regulate the atmospheric CO₂ concentration, detailed knowledge of the marine carbon cycle is needed. The well-known BCP is one of three major mechanisms by which the ocean takes up atmospheric CO₂ (Liu et al., 2018). The BCP is described by these major processes: Phytoplankton convert CO₂ into fixed carbon, e.g., carbohydrates or calcium carbonate through photosynthesis in the euphotic zone; Part of the CO₂ fixed is transferred to the ocean interior of the ocean, mainly by gravitational sinking of particulate organic carbon (POC); Diffusion, advection and vertical mixing of dissolved organic carbon (DOC) and active bio-transport of organic and inorganic carbon by diel vertical migrating zooplankton are also important carbon removal pathways; Finally very few percentage of fixed carbon will be sequestered in the deep ocean (Buesseler et al., 2007). All these biologically mediated processes constitute the BCP. Thus, sinking particles play an important role in driving the BCP (Falkowski et al., 1998; Ducklow et al., 2001; Wei et al., 2011; Liu et al., 2018).

The ratio of export production (or the POC export flux) to total primary production, known as the “f-ratio” is defined as the export efficiency of the BCP and is used to quantity the strength of the pump (Buesseler and Boyd, 2009). Hence, an investigation of the sinking POC export flux from the upper ocean would provide fundamental parameters for predicting future changes in the marine carbon cycle. To constrain the POC export flux from the upper ocean, two traditional approaches are frequently applied: sediment traps (Honjo et al., 2008; Hayes et al., 2018) and the disequilibrium of the natural radionuclide pair, ²³⁴Th-²³⁸U (Charette et al., 1999; Benitez-Nelson et al., 2001; Buesseler et al., 2009, 2020; Zhou et al., 2020).

In recent years, another pair of natural radionuclide tracers, ²¹⁰Po (T_1/2 = 138.4 days) and its progenitor ²¹⁰Pb (T_1/2 = 22.4 year), has been shown to be effective at predicting the POC export flux (Murray et al., 2005; Stewart et al., 2007, 2010; Verdeny et al., 2009; Wei et al., 2011; Le Moigne et al., 2013; Roca-Martí et al., 2016; Hayes et al., 2018; Subha Anand et al., 2018; Tang et al., 2019; Horowitz et al., 2020); however, this tracer pair has not been applied frequently. The ²¹⁰Po/²¹⁰Pb ratio exhibits different behavior from the ²³⁴Th/²³⁸U ratio in terms of particle-binding properties in the ocean. ²³⁴Th is particle reactive and is scavenged from surface waters when particles sink into the deep sea, while naturally occurring ²³⁸U is conserved and remains dissolved in well-oxygenated seawater. In contrast, both ²¹⁰Po and ²¹⁰Pb are particle reactive. In addition, ²¹⁰Pb and ²³⁴Th are only adsorbed onto the particle surface, whereas ²¹⁰Po is both adsorbed onto surfaces and biologically assimilated into cells by some species of phytoplankton and bacteria (Fisher et al., 1983; Cherrier et al., 1995; LaRock et al., 1996; Stewart and Fisher, 2003a,b). As a result, phytoplankton cells accumulating ²¹⁰Po are further ingested by zooplankton, while sinking particles excreted as fecal pellets and biogenic detritus are depleted in ²¹⁰Po (Heyraud et al., 1976). This biogeochemical behavior leads to the distribution coefficient of ²¹⁰Po being higher than that of ²¹⁰Pb (Tang et al., 2017; Zhong et al., 2019). Thus, when particles sink from the upper ocean, the large difference in particle affinity leads to secular disequilibrium between ²¹⁰Po and ²¹⁰Pb, which can be used to estimate particulate material export in a manner similar to the use of ²³⁴Th-²³⁸U disequilibrium (Friedrich and Rutgers van der Loeff, 2002; Cochran and Masqué, 2003; Verdeny et al., 2009; Wei et al., 2011; Hayes et al., 2018; Subha Anand et al., 2018; Horowitz et al., 2020). However, relatively few studies have used the ²¹⁰Po-²¹⁰Pb disequilibrium method to quantify the POC export flux at a basin wide scale (Ceballos-Romero et al., 2016; Roca-Martí et al., 2016; Tang and Stewart, 2019; Tang et al., 2019).

Oligotrophic open waters compose a major proportion (∼75%) of the surface ocean and account for over 30% of global marine carbon fixation (Shih et al., 2015). The NPO is an important atmospheric carbon sink, it can contribute to ∼25% of the total ocean CO₂ uptake per year (Takahashi et al., 2009). By assessing the carbon export in the NPO, the goal of peak carbon dioxide emissions and carbon neutrality for humanity can be achieved more scientifically, and the pace of global climate change can be gradually showed down. As a part of the “Marine Environment Monitoring and Early Warning System Construction in the West Pacific Ocean (MEMEWSC)” project, we took advantage of a unique opportunity to obtain measurements of ²¹⁰Po and ²¹⁰Pb activity over a large tract of the w-NPO (from 117°E to 146°E and from 20°N to 40°N). In this study, we describe the activity profiles (1,000 m) of ²¹⁰Po and ²¹⁰Pb at 7 stations spanning different oceanographic regions with different aeolian inputs, distances from the nearest coast and rates of primary production. The goals of this study are listed as follows: (1) To investigate the spatial variability of particle scavenging and removal processes in the water column of each oceanographic region; and (2) to estimate the magnitude and efficiency of the POC export flux derived from the ²¹⁰Po-²¹⁰Pb disequilibrium in the euphotic layer of the w-NPO.

Materials and Methods

Sampling and Preparation

The study area located in the western part of the NPO, covering an area of 20–40°N and 118–153°E. The main surface ocean currents in this region are shown in Figure 1. The Kuroshio Current (KC) is the most important western boundary current in the w-NPO, it originates from the North Equatorial Current and subsequently intrudes from the western Philippine Sea into the northern South China Sea (SCS) in the Luzon Strait. The main KC enters the East China Sea (ECS) through a channel in east of Taiwan Island and then passes through Tokara Strait at around 30°N. After the current has passed the Shikoku Basin, the northeastward KC merges with the southwestward subarctic Oyashio Current (OC) before leaving the coast of Japan to form the KC Extension in the NPO, resulting in a system known as the Oyashio-Kuroshio boundary region (Figure 1).

The MEMEWSC cruise was carried out in May-June 2018 on board RV Xiangyanghong III from the SCS to the Oyashio-Kuroshio mixing region of the western subarctic NPO (Figure 1). Dissolved and particulate samples for ²¹⁰Po and ²¹⁰Pb activity analysis were collected from the water column at 7 stations during the ∼1 month duration of the cruise. Discrete seawater samples were collected using a stainless steel sampling rosette equipped with 24 × 12 L Niskin bottles and a CTD Seabird sensor package. Seawater samples (∼20 L each) were taken from 9 depths from 0 to 1,000 m, with higher resolution in the upper 200 m of the water column. Salinity and temperature were obtained from these sensors (Seabird SBE 9/11plus CTD deployed with two Sea-Bird SBE 3P temperature sensors and two Sea-Bird SBE 4C conductivity sensors). To evaluate the POC export flux, 150 m is chosen as the integration depth according to estimated depth for the base of the euphotic zone on the ship from the Chl-a profiles.

²¹⁰Po and ²¹⁰Pb Determination

Analysis for ²¹⁰Pb and ²¹⁰Po was described in Zhong et al. (2019). In brief, seawater samples (18–24 L each) were rapidly filtered through 142-mm diameter Nucleopore filters (0.45 μm pore size) to separate the particulate and dissolved ²¹⁰Po and ²¹⁰Pb. The filtered (dissolved) seawater was then acidified to pH 1–2 with concentrated HCl immediately after collection and spiked with a known amount of ²⁰⁹Po (No. 7299, Eckert and Ziegler Isotope Products) and stable Pb²⁺ to quantify any subsequent losses of Po and Pb. After equilibration for 6–12 h, 100 mg of Fe³⁺ was added to the solution, the pH was adjusted to approximately 8–9, using concentrated NH₄OH to co-precipitate Po and Pb with Fe(OH)₃ precipitate. After settling for 8–12 h, the precipitate was transferred into a 1.5-L polyethylene bottle, and stored on board for processing upon arrival on land. In the laboratory, the precipitate was centrifuged and dissolved in 6 M HCl solution in a clean Teflon beaker, and the pH was neutralized to 1–2. Both ²¹⁰Po and ²⁰⁹Po were auto-plated onto a nickel disc (Zhong et al., 2020) after adding 0.3 g of ascorbic acid, 1 mL of 25% sodium citrate, and 1 mL of 20% hydroxylamine hydrochloride to the solution. The particulate filters were spiked with known quantities of ²⁰⁹Po and stable Pb²⁺ and totally dissolved in a mixture of HF, HNO₃ and HClO₄. After evaporating to nearly dryness several times to ensure that all the radionuclides were in dissolved form, the residue was then picked up in dilute HCl (0.1 M) for plating of Po onto a nickel disc. The plating procedure was similar to that utilized for dissolved samples. ²¹⁰Po and ²⁰⁹Po activities were determined by alpha spectrometry (Canberra series 7200-08).

For ²¹⁰Pb determination, any remaining Po isotopes in the plating solution was removed by resuspending another nickel disc, and the purified sample solution was re-spiked with additional ²⁰⁹Po and stored for ∼12 months to allow ingrowth of ²¹⁰Po from ²¹⁰Pb, and then polonium isotopes were again auto-deposited, and counted. The recoveries of ²¹⁰Pb were determined through the added stable Pb and the measured Pb, using atomic absorption spectrometry (AAS). The in-situ activities of ²¹⁰Po and ²¹⁰Pb at the sampling date were determined by correcting for decay, ingrowth, chemical recoveries, detector backgrounds, and reagent blanks (²¹⁰Pb and ²¹⁰Po present in the stable lead) (Church et al., 2012; Baskaran et al., 2013; Rigaud et al., 2013).

Particulate Organic Carbon Analysis

POC samples were obtained by filtering 4–6 L seawater through pre-combusted and pre-weighted 25 mm QMA filters (with a pore size of 0.45 μm, Whatman) immediately after sampling, and the filters were then washed 3 times with Milli-Q water to remove salt before frozen storage prior to laboratory analysis. Before the POC content was measured, the sample was oven-dried at 55°C to a constant weight. Then, the inorganic carbon was removed from the filter by using acid vapor (concentrated HCl) in desiccators for 48 h. The POC was then determined using a Vario ELIII CHNOS Elemental Analyzer. The analytical precision of the method was estimated to be around ± 4%, using triple measurement of the same sample.

Results

Hydrographic Characteristics and Particulate Organic Carbon Concentrations During the MEMEWSC Cruise

Vertical profiles of the temperature and salinity in the w-NPO are shown in Figure 2. Among the 7 stations, the sea surface temperature (SST) ranged from 15.5 to 28.9°C. The SST was high (27.6–28.9°C) at the low-latitude (21°N) stations (XTW1-3, XTW1-11 and XTJ1-1) but decreased gradually with increasing latitude from station XTJ1-1 (28.4°C) to station XTJ1-22 (15.5°C). Influenced by the OC, station XTJ1-22 showed the lowest SST. Furthermore, the water temperature decreased with increasing depth and did not differ appreciably between stations at 1,000 m (Figure 2A).

The salinity profiles spatially differed among the 7 stations (Figure 2A). At station XTJ1-22, the salinity decreased rapidly with increasing depth and exhibited a subsurface minimum (33.82) at 150 m, displaying the features of low-salinity subarctic waters (OC), and then the salinity increased gradually to 1,000 m. In contrast to station XTJ1-22, the other six stations exhibited similar salinity distributions, with the maximum values occurring at depths of approximately 150–200 m (Figure 2A), indicating the existence of stratification in the water column at each station. However, the salinity in the top 150 m of the water column at station XTW1-3 (located in Luzon Strait) was lower (from 33.80 to 34.66) than that at the station in the NPO basin (from 34.55 to 35.05) (Figure 2A). From the temperature-salinity-depth diagram (Figure 2B), the water masses can be clearly discriminated for the study area. Station XTJ1-22 showed a signature of OC, and station XTW1-3 might represent the South China Sea water. From the Figure 2B, the remaining five stations (XTJ1-1, XTW1-1, XTJ1-8, XTJ2-7, and XTJ1-13) could be divided into two groups with different hydrographic features for the surficial water mass (upper 200 m). One group includes stations XTJ1-1 and XTW1-11, and the other group contains stations XTJ1-8, XTJ1-13, and XTJ2-7, showing a much narrower ranges of salinity and temperature (Figure 2B). In 1,000 m, all the stations displayed a similar temperature-salinity feature, with low temperatures of < 5°C and a salinity range of 34.3–34.5. Based on the salinity and temperature profiles, the thermocline was observed between 100 and 200 m. To determine the euphotic zone depth, we use information for the base of the euphotic zone according to the subsurface chlorophyll-a maximums (SCMs) features. The profiles of chlorophyll-a concentrations showed that SCMs could occur between 120 and 150 m for most of the stations (33 out of 43 stations) during the cruise (unpublished data, personal communication with Dr. Jianhua Kang). Hence the euphotic zone could be delimited from 0 to 150 m. Overall, the water column could be divided into two boxes: 0–150 m box (the euphotic zone) and 150–1,000 m box (the mesopelagic zone).

All POC concentration data can be found in Supplementary Table 1. In total, the POC concentrations ranged from 1.85 to 14.35 μmol C/L, with an average of 4.15 ± 2.15 μmol C/L in the w-NPO (Figure 3). Station XTJ1-22 at the Kuroshio Current-Oyashio Current boundary had the highest POC concentration (mean: 6.23 ± 4.12 μmol C/L) in the upper 150 m, followed by station XTW1-3 (mean: 5.44 ± 1.28 μmol C/L) in the Luzon Strait, while the other five stations had similar POC concentrations with a mean value of 3.12–3.81 μmol C/L (Table 1). Almost all POC profiles showed a decrease with increasing water column depth from 0 to 500 m, and remained relatively constant from 500 to 1,000 m except at two stations (XTJ1-8 and XTJ1-13) (Figure 3).

²¹⁰Po and ²¹⁰Pb Profiles

Profiles of dissolved ²¹⁰Po (D-Po), particulate ²¹⁰Po (P-Po), total ²¹⁰Po (T-Po), dissolved ²¹⁰Pb (D-Pb), particulate ²¹⁰Pb (P-Pb) and total ²¹⁰Pb (T-Pb) were displayed in Figures 4, 5, respectively. From Figures 4, 5, overall, P-Po and P-Pb showed a similar variation trend, but D-Po showed a little bit of difference comparing with the D-Pb. The activity concentrations of particulate ²¹⁰Po (P-Po) ranged from 1.4 to 10.4 dpm/100 L (Figure 4), which are comparable to the reported values in the upper Sargasso Sea (>500 m, 0.5–8.2 dpm/100 L) (Kim and Church, 2001), and slightly higher than those in the Aleutian Basin (negligible to 6.0 dpm/100 L) (Hu et al., 2014) and in the North Atlantic Ocean (0.1–7.0 dpm/100 L) (Horowitz et al., 2020). All stations showed a systematic decrease in P-Po activity with depth in the upper 1,000 m, except for station XTW1-3 in the Luzon Strait, which exhibited a subsurface maximum between 200 and 500 m (Figure 4).

The activities of dissolved ²¹⁰Po (D-Po) varied from 2.87 to 14.57 dpm/100 L (Figure 4), which are comparable to other results previously measured in the w-NPO (5.28–14.02 dpm/100 L) (Nozaki and Tsunogai, 1976), the North Atlantic Ocean (0.6–13.1 dpm/100 L) (Horowitz et al., 2020) and the western equatorial Pacific Ocean (4.26–14.82 dpm/100 L) (Peck and Smith, 2000), although these regions might correspond to very different biogeochemical conditions. The D-Po concentration was lowest in surface seawater and gradually increased with depth; the subsurface D-Po peak was observed within approximately 50–150 m at all stations (Figure 4). This increase in the D-Po concentration with depth is consistent with the rapid scavenging of ²¹⁰Po near the surface and the remineralization of particulate ²¹⁰Po at depth. The profiles of total ²¹⁰Po (T-Po) showed vertical variations similar to those of P-Po, with a range of 4.61–20.69 dpm/100 L.

The activity concentrations of dissolved ²¹⁰Pb (D-Pb) and particulate ²¹⁰Pb (P-Pb) ranged from 1.86 to 24.45 dpm/100 L and from 1.88 to 31.33 dpm/100 L, respectively (Figure 5). The P-Pb activity was highest in the near-surface seawaters at all stations; thus, ²¹⁰Pb activity was predominantly in the particulate phase (P-Pb/T-Pb > 50%) in the upper 50 m (Figure 6). Below 50 m, P-Pb showed a systematic downward trend with depth. However, the activity of D-Pb was lowest in the surface seawater; then, D-Pb increased with depth from 50 to 200 m and finally decreased below 200 m. All stations had the highest total ²¹⁰Pb (T-Pb) concentration at the surface, and T-Pb decreased with depth from the surface to 1,000 m, suggesting a surface source of ²¹⁰Pb due to atmospheric input, as indicated by the ²¹⁰Pb activity (7.22–38.61 dpm/100 L) being greater than the ²²⁶Ra activity (6.0–14.4 dpm/100 L, Kawakami and Kusakabe, 2008) in the upper water column.

At all stations, the percentages of P-Po in the w-NPO surface seawater (0–50 m) were higher than 50%, showing that P-Po often exceeded D-Po and dominated T-Po. The high fraction of particulate ²¹⁰Pb indicated a signature of atmospheric input. Moreover, the P-Po/T-Po percentages decreased slowly to constant values of 20–30% as the water depth increased up to 500 m (Figure 6). P-Pb was generally 20–50% of the T-Pb at depths below the mixed layer, demonstrating that T-Pb was dominated by D-Pb in most layers of the water column at all stations in the w-NPO (Figure 6). Overall, these findings reveal that both ²¹⁰Po and ²¹⁰Pb were predominantly present in the dissolved phase below 50 m at all stations (Figure 6), as is commonly the case in other oceans.

²¹⁰Po/²¹⁰Pb Activity Ratio

Vertical profiles of the ²¹⁰Po/²¹⁰Pb activity ratios in the particulate, dissolved and total fractions are plotted in Figure 7. Particulate matter was enriched in ²¹⁰Po (particulate ²¹⁰Po/²¹⁰Pb activity ratios > 1) at the subsurface (50–100 m) at all stations, especially in the particulate sample from a depth of 75 m at station XTJ1-1, where the P-Po/P-Pb ratio reached 3.04 ± 0.73 (Figure 7). These data indicates that there was an excess of P-Po relative to P-Pb. The dissolved ²¹⁰Po/²¹⁰Pb activity ratios (D-Po/D-Pb) increased with depth, with values < 1 at the surface and increasing toward 1 or even > 1 below the euphotic zone, which complemented the elevated P-Po/P-Pb ratio. All stations showed similar vertical distributions of T-Po/T-Pb with a significant deficiency of ²¹⁰Po relative to ²¹⁰Pb from the surface to 200 m. The deficit of ²¹⁰Po relative to ²¹⁰Pb declined with depth and then approached equilibrium or shifted to an excess (²¹⁰Po/²¹⁰Pb≈1 or > 1) at different depths for different stations, except for stations XTJ1-1 and XTJ1-8, which displayed an excess of T-Po relative to T-Pb (T-Po/T-Pb ratio > 1) at 50–75 m and at 500–1,000 m. Interestingly, ²¹⁰Po excess (T-Po/T-Pb ratios > 1.2) was observed in seawater at 500 and 1,000 m of station XTJ1-1 and in seawater at 1,000 m of station XTJ1-8 (Figure 7). From Figure 7, we can find that this ²¹⁰Po excess is caused by the high dissolved ²¹⁰Po activity concentration in the deep layers (500–1,000 m), because D-Po/D-Pb ratios at deep layers (500–1,000 m) were much higher than unity (Figure 7). And this high dissolved ²¹⁰Po activity concentration implied the occurrence of a strong ²¹⁰Po dissolution in the deep ocean.

Discussion

Enhanced Particle Scavenging of ²¹⁰Po and ²¹⁰Pb and Deficiency of ²¹⁰Po in the Western North Pacific Ocean

Both Pb and Po are particle-reactive elements that can be scavenged effectively from the water column by lithogenic and biogenic particles. Figure 4 reveals that both the P-Po and the P-Pb activity concentrations were highest in the near-surface (0–30 m) layers at all stations; additionally, at most stations, the highest POC concentrations were similarly detected in the upper water column (>100 m), as shown in Figure 3, which resulted in higher percentages of P-Po/T-Po and P-Pb/T-Pb ratios (>50%) in the near-surface (30–50 m) seawater. The high fraction (30–60%) of particulate ²¹⁰Pb for most stations was related to atmospheric deposition of ²¹⁰Pb. In the open ocean, the most important input of ²¹⁰Pb is the atmospheric deposition, and the ²¹⁰Pb is mainly in dissolved phase in the atmospheric deposition samples. ²¹⁰Pb is particle-reactive, therefore, in the upper layer with higher primary productivity, some ²¹⁰Pb sourced from the atmosphere would be absorbed and enriched by the biogenic particles. In addition, our study area is under the cover of dust deposition from the East Asia (Hayes et al., 2013); during every spring (especially in April-May), the prevailing dust storms from the East Asia would possibly increase the particulate ²¹⁰Pb activity.

Furthermore, the P-Po/P-Pb ratios were nearly 1 or higher than 1 within 30–75 m at all stations (Figure 7), supporting the previously findings that ²¹⁰Po was preferentially scavenged over ²¹⁰Pb (Bacon et al., 1976; Murray et al., 2005; Tang and Stewart, 2019). In addition, on all profiles, both P-Po and P-Pb decreased with depth in the upper 100 m, which is consistent with the increases in D-Po and D-Pb (Figures 4, 5). The highest P-Po activity concentrations were presented at the surface for all stations (Figure 4), which could be related to particle concentration, since particle concentration is usually higher than in the twilight zone or deeper layers. D-Po increased from the surface to 1,000 m (Figure 4), especially at stations XTW1-11 and XTJ1-8, which showed subsurface D-Po maxima at 150 and 50 m, respectively. This phenomenon implies that sinking particulate matter is re-mineralized at depth in the w-NPO, which is supported by an excess of D-Po relative to ²¹⁰Pb, with the D-Po/D-Pb activity ratios often > 1 below 200 m (Figure 7). Specifically, at station XTW1-11, the feature of D-Po/D-Pb > 1 appeared between 150 and 1,000 m, while at station XTJ1-1, D-Po/D-Pb > 1 only appeared between 500 and 1,000 m. Except for station XTJ2-7, the feature of D-Po/D-Pb > 1 appeared at 1,000 m for all stations (Figure 7). Based on the above discussion, it can be concluded that enhanced particle scavenging of both ²¹⁰Po and ²¹⁰Pb occurred in the upper ocean and that ²¹⁰Po adsorbed onto particles from the upper ocean was released in the deeper ocean due to the remineralization of particulate organic matter below the mixed layer.

The minimum activity concentration of ²¹⁰Po appeared at the 100 m layer at XTW1-3, XTJ1-8 and XTJ2-7, while the minimum ²¹⁰Po activity concentration appeared at the 200 m layer at station XTW1-11, XTJ1-1, XTJ1-13, and XTJ1-22 (Figure 4). As shown in Figure 7, T-Po was largely deficient relative to T-Pb in the upper 200 m of the w-NPO, with the T-Po/T-Pb ratios as low as 0.4. This phenomenon is similar to the general pattern in which T-Po is usually deficient with respect to T-Pb (T-Po/T-Pb ratio = 0.5) in the euphotic zone due to the rapid removal of Po caused by the sinking of biogenic particles (Bacon et al., 1976; Nozaki et al., 1998; Stewart et al., 2010; Roca-Martí et al., 2016; Subha Anand et al., 2018). In contrast, ²¹⁰Po gradually reached secular equilibrium with ²¹⁰Pb or even an excess relative to ²¹⁰Pb (Figure 7) at the bottom of the euphotic zone (∼150 m) or in the mesopelagic waters (500 or 1,000 m) due to the regeneration of ²¹⁰Po from sinking particles during particle remineralization. For example, at stations XTJ1-1 and XTJ1-8, T-Po/T-Pb ratios were high up to 1.61–2.02 at 500–1,000 m layers. From Figures 4, 5, we can clearly find that the D-Po activity concentration increased significantly, but at the same layers, both D-Pb and P-Pb activity concentration decreased significantly. The explanation for this increased D-Po activity at the deep layer may be related to strong particle export events. We believe that a strong particle export event occurred in the upper ocean before our sampling, and subsequently sinking particles dissolved at the deep layers. This ultimately increased the activity concentration of dissolved ²¹⁰Po in the mesopelagic zone. Similar research cases have also been reported in a review paper published by Verdeny et al. (2009). Such disequilibrium between ²¹⁰Po and ²¹⁰Pb in the water column could help depict the scavenging rates of ²¹⁰Po and ²¹⁰Pb and to estimate the removal flux of particulate matter in the w-NPO.

Scavenging and Removal Fluxes of ²¹⁰Po and ²¹⁰Pb in the Western North Pacific Ocean

²¹⁰Pb and ²¹⁰Po in the upper ocean come from both atmospheric deposition and in situ production via parent radionuclide decay. Generally, residence times of ²¹⁰Po and ²¹⁰Pb in the atmosphere are only several days to several weeks. Hence, ²¹⁰Po deposition fluxes to the surface ocean are only approximately 10–20% of those of ²¹⁰Pb (Masqué et al., 2002; Baskaran, 2011). Many previous researchers have applied the ²¹⁰Po/²¹⁰Pb disequilibrium method to trace the POC export fluxes in the open ocean, polar sea or remote sea areas, but the ²¹⁰Po and ²¹⁰Pb deposition fluxes have consistently been neglected due to a lack of relevant data (Friedrich and Rutgers van der Loeff, 2002; Murray et al., 2005; Le Moigne et al., 2013; Roca-Martí et al., 2016; Subha Anand et al., 2018; Tang et al., 2019; Horowitz et al., 2020). According to a compilation of global ²¹⁰Pb fallout data by Du (2019), among all the observation stations worldwide, Taiwan, Shanghai and Tatsunokuchi recorded the highest annual ²¹⁰Pb deposition fluxes, indicating that East Asia exhibits the highest ²¹⁰Pb deposition flux in the world. Thus, when applying ²¹⁰Po-²¹⁰Pb disequilibrium to study particle dynamics and to estimate the POC export fluxes in the seas of East Asia, the atmospheric inputs of ²¹⁰Po and ²¹⁰Pb and their impacts on the POC export flux should be considered.

The ²¹⁰Po and ²¹⁰Pb activities in the ocean are the result of a balance among atmospheric inputs, continuous production from the decay of mother nuclides (²¹⁰Pb or ²²⁶Ra) in seawater, the radioactive decay of ²¹⁰Po and ²¹⁰Pb, removal onto sinking particles, and transport into or out of the system by advection and diffusion. Generally, the general form of the mass balance equation in the upper ocean for ²¹⁰Po and ²¹⁰Pb between sources and sinks can be designed as follows:

where ∂⁡Pb/∂⁡t and ∂⁡Pb/∂⁡t are the changes in the ²¹⁰Pb and ²¹⁰Po activities with time, respectively; F_Pb and F_Po (dpm/m²/d) are the atmospheric deposition fluxes of ²¹⁰Pb and ²¹⁰Po to the sea surface, respectively; λ_Pb and λ_Po are the decay constants of ²¹⁰Pb and ²¹⁰Po (d^–1), respectively; I_Ra, I_Pb and I_Po (dpm/m²/year) are the inventories of ²²⁶Ra, ²¹⁰Pb, and ²¹⁰Po, respectively; k_Pb and k_Po are the scavenging (from dissolved to particulate) rate constants of ²¹⁰Pb and ²¹⁰Po, respectively; and V (dpm/m²/d) is the sum of the advection and diffusion fluxes.

The advection-diffusion term is relatively important only in the case of algal blooms, mesoscale eddies, and upwelling regions (Tang et al., 2019; Horowitz et al., 2020). Generally, a steady-state model can be adopted by ignoring advection and diffusion (V≈0). Figure 8 displays the one-dimensional irreversible scavenging conceptual model for ²¹⁰Po and ²¹⁰Pb in surface and deep waters. Thus, Eqs. (1) and (2) can be rearranged as follows. For the upper ocean (0–150 m), we obtain:

where Eqs. (3) and (4) are mass-balance equations for D-Pb and P-Pb, and Eqs (5) and (6) are mass-balance equations for D-Po and P-Po.

The residence times of ²¹⁰Po and ²¹⁰Pb in the upper ocean can therefore be written as:

Similarly, for the deeper ocean (150–1,000 m), we obtain:

The residence times of ²¹⁰Po and ²¹⁰Pb in the deeper ocean can therefore be written as:

In the above equations, I_Pb and I_Po are the inventories of ²¹⁰Pb and ²¹⁰Po, respectively, in the different boxes; P_Pb and P_Po are the removal fluxes of P-Pb and P-Po, respectively, by particles sinking out of the box; J_Pb and J_Po are the rates of scavenging of ²¹⁰Pb and ²¹⁰Po, respectively (if a radionuclide is transferred from solution to the particulate phase by particle scavenging, then J is positive, while J is negative for opposite); and τ_P and τ_D are the residence times of the radionuclides in the particulate and dissolved phases, respectively.

The average annual atmospheric ²¹⁰Pb deposition flux (F_Pb) in East Asian is 18540 dpm/m²/yr (n = 25, Zhong, 2020), while the atmospheric flux of ²¹⁰Po (F_Po) is usually assumed to be only 10% of F_Pb (Turekian et al., 1977; Masqué et al., 2002; Baskaran, 2011). Table 1 shows the scavenging rates, removal fluxes and residence times of ²¹⁰Po and ²¹⁰Pb in the near-surface and deeper ocean. Inventories of ²²⁶Ra were calculated based on published data from station KH-71-3-S (which is very close to station XTJ1-8) in the w-NPO (Nozaki and Tsunogai, 1976), the ²²⁶Ra inventories in the 0–150 m and 150-1,000 m boxes were calculated to be 14,625 dpm/m² and 172,530 dpm/m², respectively (see Table 2).

Within the 0–150 m surface layer, J_Po ranged from 1,693 to 17,644 dpm/m²/yr; these positive values confirm that ²¹⁰Po was transferred from the dissolved phase to the particulate phase due to the scavenging of sinking particles in the upper ocean. Correspondingly, the removal flux of ²¹⁰Po ranged from 3,327 to 34,941 dpm/m²/year, showing a large spatial variation in the w-NPO. The residence times of D-Po (τ_Po–D) and P-Po (τ_Po–P) in the upper 0–150 m layer ranged from 0.6 to 5.5 year and from 83 to 921 days, respectively. Interestingly, the values of τ_Po–D (0.6–1.2 year) and τ_Po–P (83–93 days) at the three more northerly stations (XTJ1-13, XTJ1-22, and XTJ2-7) were much lower than the values of τ_Po–D (1.6–5.5 year) and τ_Po–P (130–921 days) of the more southerly stations (XTW1-3, XTW1-11, XTJ1-1, and XTJ1-8). Generally, the residence time of 0.6 years was common for dissolved ²¹⁰Po under more biologically productive conditions in the ocean (Shimmield et al., 1995). The τ_Po–D at stations XTJ1-13, XTJ1-22, and XTJ2-7 were much shorter than the other four stations, and correspondently, the average POC concentration at 0–150 m box were higher than the other four stations (Table 1), which supports that the residence time of dissolved ²¹⁰Po was shorter when the biological activity was higher. In addition, longer particulate residence times are related to inefficient vertical removal processes (low export fluxes), which can be confirmed from the removal flux of ²¹⁰Po (P_Po). As seen from Table 1, the P_Po values of the three more northerly stations (3,327–17,242 dpm/m²/year) were much lower than those of the four more southerly stations (14,785–34,941 dpm/m²/year). The above phenomena suggest that the upper waters of the southern part of the w-NPO are characterized by inefficient vertical removal processes and high turnover rates, while strong particle export occurred in the northern part of the w-NPO.

At greater depths (150–1,000 m), the J_Po values were negative at all stations except stations XTJ1-13 and XTJ2-7 (Table 1), further demonstrating the occurrence of ²¹⁰Po remineralization in the deeper ocean. However, due to the release of ²¹⁰Po, τ_Po–D could not be calculated at almost all stations in the deeper layer (150–1,000 m) except at station XTJ1-13. The P_Po values were negative at stations XTJ1-1 and XTJ1-8, implying that ²¹⁰Po could not be removed from the 150–1,000 m layer to the much deeper (>1,000 m) ocean at these two sampling sites. In addition, the τ_Po–P in the 150–1,000 m layer ranged from 85 to 331 d at the remaining four stations (XTW1-3, XTJ1-13, XTJ 1-22, and XTJ2-7).

In the upper 0–150 m, the scavenging rates of ²¹⁰Pb (J_Pb) varied between 18,373 and 18,720 dpm/m²/yr and the removal fluxes of ²¹⁰Pb (P_Pb) ranged from 17,835 to 18,454 dpm/m²/year (Table 1). As atmospheric input is the dominant source term for ²¹⁰Pb in the upper ocean, the removal flux of ²¹⁰Pb was close to the constant atmospheric ²¹⁰Pb deposition flux of 18,540 dpm/m²/year. The values of τ_Pb–D and τ_Pb–P in the upper ocean (0–150 m) were nearly on the same level, ranging from 0.47 to 1.09 year and from 0.33 to 0.98 year, respectively. However, in the deeper ocean (150–1,000 m), the scavenging rate of ²¹⁰Pb decreased significantly to a low level of 3,103–3,800 dpm/m²/year, whereas the removal flux of ²¹⁰Pb remained at a high level (15,575–17,649 dpm/m²/year). Moreover, the residence times of D-Pb and P-Pb below 150 m were calculated to be 13.1–22.9 and 1.1–5.5 year, respectively.

Particulate Organic Carbon Export Flux Estimated From ²¹⁰Po Deficiency

The scavenging of particle-reactive nuclides produces a deficiency of the daughter nuclides with respect to their parents in the water column. From these daughter deficiencies, Buesseler et al. (1992) developed a method to estimate the POC export flux (F_POC) from the upper ocean by using ²³⁴Th-²³⁸U disequilibrium, and this method has been applied extensively. F_POC can be derived by multiplying the daughter nuclide removal flux from the euphotic zone by the ratio of the POC concentration to the daughter nuclide activity of the total particulate material. Similarly, based on the ²¹⁰Po deficiency, the POC export flux is expressed as follows:

where $POC/A_{Po}^P$ denotes the measured POC/²¹⁰Po ratio in particulate matter, and P_Po denotes the removal flux of ²¹⁰Po at the output interface of 150 m.

Similarly, Coale and Bruland (1987) showed that the profile of ²³⁴Th was closely linked to the profiles of nutrients and chlorophyll, implying that the cycling of ²³⁴Th was linked to the cycling of organic matter. This led Eppley (1989) to assume that, if the residence times of POC and ²³⁴Th in surface waters are similar, the ²³⁴Th flux can be used to calculate F_POC. Because ²¹⁰Po has stronger particle affinity in organic matter than ²³⁴Th, ²¹⁰Po could be superior to ²³⁴Th as a tracer for POC; hence, based on Eppley’s assumption, the ²¹⁰Po-based POC flux can also be derived as follows:

where I_POC stands for the integral inventory of POC in the upper ocean and τ_Po–P is the residence time of ²¹⁰Po in the particulate phase.

The POC fluxes estimated by the Eppley and Buesseler models in the upper layer of the w-NPO (0–150 m) are displayed in Table 3 and Figure 9. From Figure 9A, we can divide these seven stations into two groups. One group (Group I) includes the station XTW1-3 in the Luzon Strait, station XTJ2-7 in the northern Shikoku Basin, and station XTJ1-22 located in the Oyashio Current -Kuroshio Current boundary near the main island of Japan, which showed a higher estimated F_POC level of 6–9 mmol C/m²/d (see Table 3 and Figure 9A). The common characteristic of Group I was that all of them were located near the continental shelf of East Asia. However, station XTJ1-22 was also affected by the cold and nutrient-rich Oyashio Current, which appears to promote high primary production and subsequently high POC export flux. Therefore, the FPOC of station XTJ1-22 was the highest among all stations (Figure 9A).

By contrast, the other four stations (XTJ1-1, XTW1-11, XTJ1-8, and XTJ1-13 station) can form another group (Group II), which denoted a lower POC export flux level in the w-NPO (Figure 9A). Concretely, the lowest POC export (<1 mmol C/m²/day) was found at station XTJ1-1 in the Marianas Trench, at which the negligible sinking flux of ²¹⁰Po (3,327 dpm/m²/year, Table 1) was the most obvious feature. The relatively lower POC export fluxes (1–2 mmol C/m²/day) at stations XTW1-11 and XTJ1-8 were derived from the relatively low sinking fluxes of ²¹⁰Po (Table 1) and relatively low particulate POC/²¹⁰Po ratios (Table 3). Geographically, these four stations of Group II (XTW1-11, XTJ1-1, XTJ1-8, and XTJ1-13) were located in the central gyre of the w-NPO, which had the lower particulate POC/²¹⁰Po ratios in total suspended matters comparing with values of the stations (XTW1-3, XTJ1-22, and XTJ2-7) in Group I close to the East Asian continent (Table 3). F_POC of stations in Group II were much lower than those of Group I, which may be related to the dominated phytoplankton. The central gyre of the w-NPO (stations of Group II) is an oligotrophic area where the phytoplankton community is dominated by picoplankton (such as smaller diatoms or coccolithophores), and these small size phytoplankton produce a low POC flux due to the oligotrophic living conditions in this region. Interestingly, in these areas, stations in Group II had much higher particulate POC/²¹⁰Po ratios than those of stations in the central gyre of the w-NPO (Group I). For example, the stations XTW1-3, XTJ1-22, and XTJ2-7, located at the margins of the Chinese mainland and the main island of Japan, have higher particulate POC/²¹⁰Po ratios of 181 ± 16, 184 ± 16, and 108 ± 10 μmol/dpm, respectively. And the other four stations XTW1-11, XTJ1-1, XTJ1-8, and XTJ1-13, located in the central gyre of the w-NPO, have much lower particulate POC/²¹⁰Po ratios of 73 ± 5, 76 ± 6, 84 ± 7, and 69 ± 5 μmol/dpm, respectively (Table 2). Generally, smaller phytoplankton cells can scavenge more ²¹⁰Po (smaller particles dominate higher particulate ²¹⁰Po activity relative to larger particles in the non-oligotrophic area) due to their larger surface area per unit of volume, lowering their particulate POC/²¹⁰Po ratios (Tang et al., 2019).

Overall, the POC fluxes calculated by the Eppley method and Buesseler method ranged from 0.6 to 8.8 mmol C/m²/d and from 0.7 to 8.6 mmol C/m²/d, respectively (Figure 9A). These values are in good agreement with the estimated POC fluxes derived via the steady-state ²¹⁰Po-²¹⁰Pb method in other open oceans with similar latitudes (negligible to 8.5 mmol C/m²/day) (Kim and Church, 2001; Stewart et al., 2007; Verdeny et al., 2009; Roca-Martí et al., 2016; Tang et al., 2019; Horowitz et al., 2020). From Figure 9A, although there are some differences, Eppley model and Buesseler model-derived F_POC, are very similar within uncertainties. Buesseler model estimated F_POC were 24 and 43% higher than that of Eppley model for stations XTW1-3 and XTJ2-7, respectively. However, Eppley model-derived F_POC was 17% higher than that of Buesseler model for station XTJ1-22. The difference of F_POC results given by different models may be related to the assumptions of the models. For example, the Eppley model’s assumption that the residence times of POC and particulate daughter radionuclides (²³⁴Th or ²¹⁰Po) should be the same within the upper ocean is possibly too demanding in the real ocean and therefore is often difficult to achieve. Murray et al. (1989) compared the residence times of ²³⁴Th and POC in the euphotic zone of the eastern equatorial Pacific and found that the residence time of ²³⁴Th was approximately three times less than that of POC. In contrast, the Buesseler model needs to assume only that the same particulate matter serves as the carrier of POC and ²³⁴Th (or ²¹⁰Po), which is relatively easy to achieve in the actual sea. For some specific stations, the difference between two models may be significant, hence, more care should be taken when selecting the model. But on a larger spatial scale in our study area, Eppley model can also provide reasonable estimation results. For example, Figure 9B shows that Eppley model-derived F_POC was well correlated with the Buesseler model-derived F_POC in the w-NPO (n = 7, slope = 1.06, R² = 0.92, p = 0.003); additionally, the ²¹⁰Po-POC export fluxes based on the Eppley and Buesseler models had very similar spatial distribution features (Figure 9A). These evidences suggest that the hypothesis of the Eppley model is also valid in the w-NPO and this agreement of Eppley model-derived F_POC and Buesseler model-derived F_POC in the w-NPO probably suggest the residence times of ²¹⁰Po and POC were similar in the upper ocean.

Comparison of Particulate Organic Carbon Export Fluxes Among Different Regions of the Western North Pacific Ocean and Its Surrounding Marginal Seas

The ²¹⁰Po-derived POC export fluxes from the upper ocean (5.9–8.8 mmol C/m²/day) at the stations along the edges of the w-NPO (close to the East Asian land masses) were significantly higher (by a factor of 2∼14) than the fluxes determined from the interior (oligotrophic) seawater in the w-NPO (i.e., XTW1-11, XTJ1-1, and XTJ1-8: 0.6–2.5 mmol C/m²/d) (Table 3 and Figure 9). Comparing our POC fluxes with other fluxes in different regions of the Pacific Ocean and its surrounding marginal seas (Table 4, in which the POC fluxes were estimated either by sediment traps or by the ²³⁴Th-²³⁸U and ²¹⁰Po-²¹⁰Pb disequilibrium methods), reveals systematic higher POC fluxes occurring in the marginal seas surrounding the w-NPO (Figure 10). For example, the POC export fluxes in the Pacific Ocean were much smaller than those reported for the SCS (2.0–16.7 mmol C/m²/day, ²³⁴Th-²³⁸U disequilibrium, Zhou et al., 2020), inner shelf of the ECS (4.8–65.4 mmol C/m²/day, trap, Hung et al., 2013), and the Sea of Japan (3.2–26.3 mmol C/m²/day, ²³⁴Th-²³⁸U disequilibrium, Kawakami et al., 2015).

The NPO is a region of particular interest for quantifying the rate and efficiency of the BCP, as this region is a major sink for atmospheric CO₂, absorbing 0.5 Pg C/year, ∼25% of the total ocean CO₂ uptake (Takahashi et al., 2009). As an important part of the NPO, the subtropical oligotrophic w-NPO is a typical nitrate-deficient water region and is characterized by low chlorophyll-a (Chl a) concentrations (Chen et al., 2013), and this region occurs a band of strong CO₂ uptake between 30 and 40°N, overlapping with the transition zone between the subtropical and subarctic gyres (Shih et al., 2015). Researchers stated the importance of the contributions of biological carbon export to the NPO sink (Ayers and Lozier, 2012). However, the characteristics of the POC export flux in the w-NPO region are not well understood (Kawakami et al., 2015; Shih et al., 2015). Using the information available, based on the summary of the published POC export fluxes in the North Pacific Ocean (Table 4 and Figure 10), we can obtain a preliminary general view for the distribution feature of POC export flux. Although there were spatial differences, the overall mean level of POC export flux was below 10 mmol C/m^2/day. Specifically, the POC fluxes in our study area of the w-NPO (0.6–8.8 mmol C/m²/day) are comparable to the results of other published studies obtained with the ²³⁴Th-²³⁸U method in other Pacific Ocean regions under general production condition (Figure 10A), such as the subarctic northeastern Pacific Ocean (2.8–7.6 mmol C/m²/d, from February 1996 to February 1997, Charette et al., 1999; 2.01 ± 0.56 (n = 61) mmol C/m²/d in August/September 2018 with no bloom, Buesseler et al., 2020), station ALOHA in the North Pacific Subtropical Gyre (0.4–4.0 mmol C/m²/d, from April 1999 to March 2000, Benitez-Nelson et al., 2001; neiligible-7.0 mmol C/m²/day, in June/July 2004 with no bloom, Buesseler et al., 2009), station K2 in the western subarctic North Pacific (1.9–10.1 mmol C/m²/day, in July/August 2005 with no bloom, Buesseler et al., 2009) and the North Pacific tropical and subtropical gyre (0.1–6.3 mmol C/m²/day, in August/September 2015 with no bloom, Umhau et al., 2019).

Export efficiency can be defined as the ratio of POC export to primary production (the e-ratio, Buesseler, 1998). Although primary productivity levels were not available for the studied sampling stations, Palevsky et al. (2016) summarized the annual and seasonal net primary productivity (NPP) estimates for the western North Pacific. Since our sampling activities took place from late May to late June 2018, we preliminarily evaluated the e-ratios of the BCP in the w-NPO by using the summer NPP estimates for the regions of the North Pacific (62 ± 19 mmol C/m²/day), hence, the e-ratio values were calculated to be in the range of 1.0–14.2% (Table 3), which are similar to the previously reported values in the North Pacific (Table 4 and Figure 10B; Charette et al., 1999; Benitez-Nelson et al., 2001; Buesseler et al., 2009, 2020). As can be seen from Figure 10B, the e-ratios of biological carbon pump were more variable in the marginal seas than that in the Pacific Ocean. For example, the e-ratios were very consistent (in a level of < 20%) in the NE Pacific, North Pacific, and w-NPO. However, in the marginal seas, the POC fluxes and e-ratios changed over a very wide range (Figure 10B). One extreme case reported in the inner shelf of ECS showed that the e-ratios ranged from 10 to 90% (Hung et al., 2013). The e-ratios in our study were below 15%, suggesting a moderate BCP efficiency in the w-NPO. By comparison, it can be found that the BCP strengths in some marginal seas of the Pacific Ocean [for example, e-ratio = 9–91% in the inner shelf of ECS (Hung et al., 2013), 13–28% in the Southern ECS (Shih et al., 2013), and 7–56% in the Sea of Japan (Kim et al., 2011)] is stronger than that in the interior of the Pacific Ocean (for example, e-ratio = ∼10% at station XTJ2-7 in the Shikoku Basin, 3–4% at station XTW1-11 in the Philippine Basin, and ∼1% at station XTJ1-1 in the Mariana Trench).

We compiled a table for comparing the export efficiencies of BCP in other open seas, as shown in the Table 5. It is worth noting that the e-ratios are similar in such different areas. For example, the e-ratios < 15% observed in our study are similar to those found in the North Atlantic (Iberian Basin, 0.5–2.5%; Irminger Basin, 3 ± 3%; western European Basin, 5 ± 5%; Iceland Basin, 6 ± 6%; and Labrador Basin, 10 ± 3%, see Tang et al., 2019). The similarity between our study and Tang et al. (2019) is that both sampling activities were carried out in May-June, however, the difference is that our study area was in an oligotrophic subtropical region (medium-low latitudes) and the study area of Tang et al. (2019) was in a temperate system (medium-high latitudes). In addition, Tang et al. (2019) started sampling when the bloom was starting, hence, the POC fluxes and e-ratios were still low. Ceballos-Romero et al. (2016) pointed that export efficiency was significantly discrepant in different stages of the bloom and the export efficiency also showed a strong seasonal variability in the North Atlantic. By comparison with study case of Ceballos-Romero et al. (2016) in the North Atlantic Ocean, we can find that algal blooming has a significant influence on BCP efficiency. Ceballos-Romero et al. (2016) reported two sampling campaigns: one sampling was carried out in the pre-bloom stage (April-May 2010), and the export efficiency was high at that moment (∼50%); another sampling was carried out in the decline of the bloom and post-bloom stage (July–August 2010), and the export efficiency was very low (3%). Obviously, our study displayed results from medium-low latitudes with no bloom, therefore, the low e-ratios of 1.0–14.2% were reasonable.

In addition, the e-ratios in the w-NPO (1.0–14.2%, this study) are close to or slightly higher than that in the low-latitude tropical seas and oceans, for example, the equatorial Pacific (1–10%, Buesseler, 1998) and the Arabian Sea (1–10%, Buesseler, 1998; 0.2–11.7%, Subha Anand et al., 2017). This low e-ratios of < 10% might represent the values of BCP efficiency under general condition for tropical ocean, considering the weak seasonal variation of tropical oceans and the oligotrophic features. However, the e-ratios in the w-NPO are much lower than those reported at high-latitude sites (>25%), such as the Eurasian Basin of the Arctic Ocean (60 ± 40%, Roca-Martí et al., 2016), the Bellingshausen Sea (37% at station K, Shimmield et al., 1995) and the Weddell Sea and the Antarctic Polar Front (16–100%, Rutgers van der Loeff et al., 1997), since most of these study cases in the high-latitude seas were related to the bloom situation or the increase of chlorophyll-a, POC and plankton biomass concentrations (Shimmield et al., 1995; Rutgers van der Loeff et al., 1997). The lower export efficiencies observed in the w-NPO may be consistent with the predominance of smaller phytoplankton in the oligotrophic seawater. The oligotrophic feature might weaken the export strength of the BCP in the w-NPO. Indeed, small particles are usually slow-sinking particles that are more likely to be degraded during their descent (Villa-Alfageme et al., 2016), leading to lower export efficiencies. In our study, we also observed negative scavenging rates (J) of ²¹⁰Po (negative J values mean that radionuclides are transferred from particles to the dissolved phase) below 150 m in the w-NPO (Table 1), which supports the occurrence of particle degradation, leading to a low export efficiency. For example, the lowest J_Po values in the 150–1,000 m segment were observed at stations XTW1-11 (J_Po = –86,511 dpm/m²/year), XTJ1-1 (J_Po = –64,036 dpm/m²/year) and XTJ1-8 (J_Po = –26,022 dpm/m²/year), and correspondingly, these three stations have the lowest e-ratios (based on Eppley model: 1.0–2.7%; based on Buesseler model: 1.1–4.0% Table 3).

Conclusion

In this study, we reported the vertical distributions of D-Po, P-Po, D-Pb, and P-Pb activities in the w-NPO during late spring and early summer to constrain the particle dynamics (scavenging rates and removal fluxes of ²¹⁰Po and ²¹⁰Pb), and to estimate the carbon export production (POC export flux). More than 50% of the radionuclides were found in the dissolved phase below a depth of 50 m, while a small proportion was associated with the particulate phase. However, the percentages of P-Po and P-Pb were higher than 50% (even up to 90%) in the surface seawater (0–50 m). ²¹⁰Po deficits relative to ²¹⁰Pb were observed in the upper 150 m at all the stations, however such ²¹⁰Po deficiencies can even extend to deeper ocean (such as 500 m) at some stations. In the interior of the w-NPO, the excess ²¹⁰Po (total ²¹⁰Po/²¹⁰Pb > 1) activities in the deeper ocean (500–1,000 m) at stations XTJ1-1 and XTJ1-8 were attributed to the release of ²¹⁰Po due to biogenic particle release during sinking. Based on a conceptual one-dimensional irreversible scavenging model, the residence times of ²¹⁰Po and ²¹⁰Pb in the w-NPO were obtained. In the upper ocean (0–150 m), the residence times of D-Po (τ_Po–D, 0.6–5.5 year) were much higher than those of P-Po (τ_Po–P, 83–921 days), while in the deep ocean (below 150 m), the residence times of D-Pb were calculated to be 13.1–22.9 year, 2–10 times longer than those of P-Pb (1.1–5.5 year).

Based on ²¹⁰Po-²¹⁰Pb disequilibrium, we found that the Eppley model-derived POC export fluxes agreed well with those derived from the Buesseler model in the upper ocean of the w-NPO, suggesting that particulate ²¹⁰Po and POC may have similar residence time in the water column. Overall, the ²¹⁰Po-derived POC fluxes varied spatially, ranging from 0.6 to < 9 mmol C/m²/day, with the highest export fluxes at stations close to the East Asian continental shelf and the lowest export fluxes at stations in the ocean basin. We concluded that POC export fluxes tended to increase with decreasing distance from the continental margins. The POC export efficiencies also showed regional differences even within the same basin, with e-ratio values ranging from 1.0 to 14.2%, suggesting a moderate BCP efficiency in the w-NPO. The low export efficiencies may be associated with the dominance of smaller particles and particle release below the euphotic zone. The negative scavenging rates of ²¹⁰Po (-J_Po) below 150 m at stations in the central basin of the w-NPO supports the occurrence of particle dissolution. Comparing with the reported e-ratios in other open seas, relatively higher carbon export efficiencies occurred at high-latitude sites.

Data Availability Statement

The original contributions presented in the study are included in the article/Supplementary Material, further inquiries can be directed to the corresponding author/s.

Author Contributions

QZ was responsible for conceptualization, methodology, sample analysis, data processing, writing original draft, writing–review, and revising. DH, TY, and JD provided a financial support and helped in article reviewing. JL, JJ, and JN contributed to the review of the manuscript. All authors contributed to the article and approved the submitted version.

Funding

This study was funded by the China Postdoctoral Science Foundation (2021M693780), the Scientific Research Foundation of Third Institute of Oceanography, MNR (2019004), the Science and Technology Plan Projects of Guangxi Zhuang Autonomous Region (2017AB30024) and the Science Foundation of the Fujian Province (2018Y0059 and 2019Y0073).

Conflict of Interest

The authors declare that the research was conducted in the absence of any commercial or financial relationships that could be construed as a potential conflict of interest.

Publisher’s Note

All claims expressed in this article are solely those of the authors and do not necessarily represent those of their affiliated organizations, or those of the publisher, the editors and the reviewers. Any product that may be evaluated in this article, or claim that may be made by its manufacturer, is not guaranteed or endorsed by the publisher.

Acknowledgments

We are grateful to the captain and crew of the R/V Xiangyanghong III for their support at sea. Many thanks to members of RIC group in ECNU (especially Lady Langlin) for their assistance with expedition preparation.

Supplementary Material

The Supplementary Material for this article can be found online at: https://www.frontiersin.org/articles/10.3389/fmars.2021.700524/full#supplementary-material

References