Reconstructing organic matter sources and rain rates in the southern West Pacific Warm Pool during the transition from the deglaciation period to early Holocene

Chemical Geology 529 (2019) 119291

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Reconstructing organic matter sources and rain rates in the southern West Pacific Warm Pool during the transition from the deglaciation period to early Holocene

T

⁎

Linying Chena, Min Luoa,b,e, , Andrew W. Dalec, Harunur Rashida, Gang Lina, Duofu Chena,d a

Shanghai Engineering Research Center of Hadal Science and Technology, College of Marine Sciences, Shanghai Ocean University, Shanghai 201306, China Laboratory for Marine Geology, Qingdao National Laboratory for Marine Science and Technology, Qingdao 266061, China c GEOMAR Helmholtz Centre for Ocean Research, Kiel 24148, Germany d Laboratory for Marine Mineral Resources, Qingdao National Laboratory for Marine Science and Technology, Qingdao 266071, China e College of Earth Science and Engineering, Shandong University of Science and Technology, Qingdao, 266510, China b

A R T I C LE I N FO

A B S T R A C T

Editor: Michael E. Boettcher

Transient features in the organic carbon content of deep-sea sediment cores resulting from changes in the flux and/or quality of sedimentary organic matter are not uncommon. We examined the geochemical characteristics of sediments retrieved with a gravity corer from the northwestern Solomon Sea (3908 m water depth), southern West Pacific Warm Pool (WPWP). δ13C and δ15N of sedimentary organic matter, together with TOC/TN data suggest that the organic material is characterized by a mixture of marine and terrestrial origins with a higher contribution from marine algae. The data were analyzed with an inverse non-steady-state reaction-transport model to examine the variability and magnitude of particulate organic carbon (POC) flux to the seafloor during the transition between the deglaciation period and early Holocene. Measured POC content and porewater NO3−, NH4+, DIC and SO42− concentrations were used to constrain the model. Hindcast results revealed that POC flux decreased from 75 to 37.5 μg cm−2 yr−1 during the deglaciation–early Holocene transition. The rate of POC degradation in the present setting is slightly lower compared to that in the pre-Holocene setting. The synchronous decline in the relative contribution of terrestrial organic matter input and rapid sea level rise during this transition suggest that sea level, rather than surface productivity, is the dominant factor controlling the POC deposition flux in the Solomon Sea. This is conceivable because the sampling site is proximal to high-relief islands with high rainfall, a well-developed submarine canyon system and narrow and steep continental margins. Consequently, we suggest that deep-water basins in proximity to similar high-relief mountainous islands in the tropical Pacific may represent important sinks for terrestrial organic material, especially during sea level lowstands.

Keywords: Particulate organic matter Transient state Reaction-transport model West Pacific Warm Pool Organic carbon burial

1. Introduction The West Pacific Warm Pool (WPWP) often refers to the equatorial western Pacific that is characterized by a sea surface temperature (SST, > 28 °C) ~2–5 °C higher than in any other equatorial region (Yan et al., 1992). Due to its large heat storage capacity, it represents one of the main sources of moisture and heat for extra-tropical regions, and may have a profound effect on global climate (Meyers et al., 1986). The WPWP is also strongly influenced by large-scale climatic features, e.g., the El Niño Southern Oscillation (ENSO) and the East Asian monsoon (Beaufort et al., 2001; de Garidel-Thoron et al., 2001; Stott et al., 2002;

Wang et al., 2004). An intense low-pressure center in the WPWP associated with the warm surface water leads to high precipitation (Partin et al., 2007). In addition, owing to its great topographic relief and active tectonics, this region is responsible for transportation of large amounts of terrestrial material to the coastal and deep ocean (Nittrouer et al., 1994; Walsh and Nittrouer, 2003; Brunskill, 2004; Wu et al., 2013). In the WPWP area, SST, upwelling, mixing of surface water, and riverine inputs are the primary factors controlling biogeochemical conditions in the surface ocean that greatly affect both the composition of planktonic communities and sinking organic detritus (Kawahata, 1999; Kawahata et al., 2002).

⁎ Corresponding author at: Shanghai Engineering Research Center of Hadal Science and Technology, College of Marine Sciences, Shanghai Ocean University, Shanghai 201306, China. E-mail address: [email protected] (M. Luo).

https://doi.org/10.1016/j.chemgeo.2019.119291 Received 9 January 2019; Received in revised form 3 June 2019; Accepted 29 August 2019 Available online 30 August 2019 0009-2541/ © 2019 Elsevier B.V. All rights reserved.

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is the Markham River and its associated submarine canyon and channel in the western Solomon Sea (Fig. 1). Sediments are supplied down the Markham Canyon and Channel as far as to the western New Britain Trench (Whitmore et al., 1999).

The vast majority of organic carbon produced in the euphotic zone through primary production is remineralized by metabolic processes within the water column and only a tiny fraction (< 1%) of organic carbon from primary production reaches the deep ocean (e.g., Sarmiento and Gruber, 2006). In spite of this, organic carbon burial in marine sediments is a key component of the Earth's carbon cycle because it provides the energy that fuels benthic ecosystems and helps to regulate atmospheric oxygen and carbon dioxide levels over geological timescales (Berner and Canfield, 1989; Bjerrum and Canfield, 2004). In addition, the content of organic carbon in the sediments partly reflects primary production and has been used as a proxy for paleomarine productivity (Schoepfer et al., 2015). The organic carbon flux to the seafloor depends largely on the biologically mediated export of carbon from the surface ocean and its remineralization with water depth. Although organic carbon flux generally decreases with depth in the form of power law (Martin et al., 1987), it shows salient spatial and temporal features on a global scale. For instance, the fraction of exported organic material from the euphotic zone reaching the deep ocean is generally higher at low latitudes than that at high latitudes due to differences in primary production and availability of biogenic minerals acting as ballast (Francois et al., 2002; Klaas and Archer, 2002). Temporally, organic carbon burial in deep-sea sediments exhibits prominent glacialinterglacial cycles with highest accumulation rates during glacial maxima (Pedersen, 1983; Bradtmiller et al., 2010). This is indicative of enhanced nutrient supply, more efficient transfer of organic matter to the deep ocean and better preservation of organic matter during glacial periods (Olivier et al., 2016). Although SST showed a small difference of ~2 °C between the present and the last glacial period (Broecker, 1986; Thunell et al., 1994; Wang, 1999), the mass accumulation rate of organic carbon in the western WPWP has fluctuated in response to glacial-interglacial alternations (Kawahata, 1999). Reaction-transport modeling represents a useful tool for investigating sedimentary organic matter fluxes and microbial activity over geological timescales (e.g., Arndt et al., 2009; Dale et al., 2012). Fluctuations in organic carbon content in deep-sea sediments generally follow glacial-interglacial cycles (Pedersen, 1983; Bradtmiller et al., 2010; Olivier et al., 2016). However, our knowledge of the deposition of particulate organic carbon (POC) during the deglaciation-early Holocene transition remains scarce. In this study, we examine the sources, deposition, and diagenetic cycling of organic material in the deep-sea sediments of the southern WPWP. Our prime objective is 1) to reconstruct the change in organic carbon deposition during the transition from the deglaciation period to early Holocene using a non-steady-state model approach and 2) to determine the driving mechanism for the change in POC flux. We suggest that sea level variation in proximity to high-relief islands in the WPWP plays a significant role in POC deposition and burial.

3. Material and methods 3.1. Sampling One gravity core (GC02) and one box core (BC01) were collected from the same site (6°36.894′S, 149°45.462′E, water depth 3908 m) located on the overriding plate of New Britain subduction zone during the RV Zhangjian expedition to the New Britain Trench in August 2016. The recovered core length was 146 cm for GC02 and 21 cm for BC01. Upon recovery, the gravity core was sealed, labeled, and split open lengthwise in the onboard lab. The porewater was collected with Rhizon samplers that were pushed into the working half of the gravity core sediments. With this technique, the porewater was extracted by suction using 20 ml plastic syringe for ~2 h with limited air contact after discarding the first milliliter. Aliquots for shore-based analysis were preserved with 10 μl saturated HgCl2 solution and stored at −20 °C. The box core sediments were subsampled for the determination of porosity using a push core immediately after retrieval.

3.2. Geochemical analysis Nitrate (NO3−) and ammonium (NH4+) concentrations in the gravity core were determined using a QuAAtro autoanalyzer (Seal Analytical) with a detection limit of 1 μM and a precision of 2%. Sulfate (SO42−) was measured by a Dionex ICS-5000+ ion chromatograph with a detection limit of 10 μM and a precision of 2%. Concentrations of dissolved inorganic carbon (DIC) and carbon isotopic compositions of DIC (δ13CDIC) were determined using a Thermo Finnigan Gas Bench connected to a Thermo Finnigan Delta V Advantage isotope ratio mass spectrometer (IRMS). Note that only DIC concentrations are shown here and δ13CDIC will be reported elsewhere. The analytical precision was better than 2% for DIC concentration. Total organic carbon (TOC), total nitrogen (TN), and carbon and nitrogen isotopic compositions of sedimentary organic matter were measured by high temperature combustion on a Vario Pyro Cube elemental analyzer connected to an Isoprime 100 continuous flow isotope ratio mass spectrometer. All samples were pre-treated with 10% HCl to remove inorganic carbon. Carbon and nitrogen isotope ratios are expressed in the delta notation (δ13C and δ15N) relative to V-PDB and atmospheric nitrogen. The average standard deviation of each measurement, determined by replicate analyses of the same sample, was ± 0.02% for TOC, ± 0.006% for TN, ± 0.2‰ for δ13C, and ± 0.3‰ for δ15N.

2. Study area The sampling site is situated on the upper plate of the western New Britain subduction zone in the southern WPWP (Fig. 1). The western New Britain subduction zone was formed by the subduction of Solomon Sea Plate beneath the South Bismark Plate (Cooper and Taylor, 1987). Hydrographically, the surface current (New Guinea Coastal Current, NGCC) in the study area is mainly driven by the trade winds, which is directed southeastward during the NW monsoon and northwestward during the SE monsoon. Below the NGCC down to ~200 m, the New Guinea Coastal Undercurrent (NGCUC) flows westward throughout the year (Kuroda, 2000). Both the NGCC and the NGCUC originate from the South Equatorial Current (SEC) and the NGCUC is the main pathway of Southern Hemisphere water into the Equatorial Undercurrent (Fine et al., 1994). The present-day water column in the WPWP is well stratified, which results in low nutrient levels in the surface and upperphotic layer due to weak vertical water mixing (Hagino et al., 2000). The main sediment conduit from Papua New Guinea to the Solomon Sea

3.3. Porosity and sediment burial velocity Porosity was determined from the weight loss before and after freeze-drying of the wet sediments. The volume fraction of porewater was calculated assuming a dry sediment density of 2.5 g cm−3 and a seawater density of 1.023 g cm−3. The sediment burial velocity, ω (cm kyr−1), was determined based on 14C-AMS ages of TOC from BC01 and GC02. The analysis was conducted at the National Ocean Sciences Accelerator Mass Spectrometry facility (NOSAMS) at Woods Hole Oceanographic Institution. 14C-AMS ages were converted to calendar years before present (yr B·P., relative to 1950 CE) using the Calib.7.1 Program with the IntCal13 calibration curve and a global average reservoir age of 400 yr (Reimer et al., 2013). Assuming steady-state compaction, ω can be determined by fitting the measured 14C ages to the following equation: 2

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Fig. 1. Schematic map of the study area showing the sampling site (yellow dot) and major modern ocean current routes. SEC = South Equatorial Current, GBRUC = Great Barrier Reef Undercurrent, CSCC = Coral Sea Coastal Current, NGCUC = New Guinea Coastal Undercurrent, NGCC = New Guinea Coastal Current, EUC = Equatorial Undercurrent. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

agex = age0 +

(Φf − Φ0 ) ∙ (e−p ∙ x − 1) x + ω p∙ω∙ (Φf − 1)

with sediment depth and was constrained with the POC, DIC, and NH4+ data. Note that a0 is independent from age0 in Eq. (1) which denotes the age of bulk organic matter at the surface sediment determined by 14CAMS dating. POM was degraded by aerobic respiration, denitrification, and sulfate reduction. As porewater profiles of Fe2+ and Mn2+ were not available, POM degradation via iron and manganese oxide reduction was neglected. In fact, their contribution to POC degradation can be considered to be negligible in deep-sea settings (Bender and Heggie, 1984). The relative rates of each POM degradation pathway were determined using Michaelis-Menten kinetics whereby the electron acceptors were used sequentially in the order of O2, NO3−, and SO42− as defined by their respective half-saturation constants (e.g., Boudreau, 1997). A bimolecular rate law was used for the aerobic oxidation of NH4. The reactions considered in the model along with kinetic rate expressions are listed in Table 1. The length of the simulated model domain was set to 200 cm. Upper boundary conditions for solutes were imposed as fixed concentrations (Dirichlet boundary) using measured values in the bottom water. They were assumed constant in time. The upper boundary for POC was prescribed as a flux (Robin-type boundary). As explained in the following section, the model was run in a non-steady-state configuration where the POC flux to the seafloor was varied over time. A zero concentration gradient (Neumann-type boundary) was imposed at the lower boundary condition for all species. Further details on the model solutions can be found in the Supplement.

(1) 14

where x (cm) is depth, agex and age0 (ka) represent the C-AMS ages of TOC at x and sediment surface, respectively, Φf and Φ0 (both dimensionless) denote the porosity below the depth of compaction and at sediment surface, respectively, and p (cm−1) is the porosity attenuation coefficient. The porosity parameters were obtained by fitting the box core data to the exponential function assuming steady-state compaction (Fig. S1). 3.4. Reaction-transport modeling A reaction-transport model was used to simulate one solid phase (POC, which is represented by TOC) and five solutes including oxygen (O2), NO3−, NH4+, SO42−, and DIC. Chemical compounds were transported by molecular diffusion (for solutes) and sediment burial (for solutes and solids). The degradation of particulate organic matter (POM) is the main biogeochemical process in the model that drives all other secondary redox reactions. POM is defined chemically as (CH2O) (NH3)rNC, where rNC is the atomic N:C ratio determined from the average of measured TN/TOC profile. The total rate of POC degradation is defined according to the power law of Middelburg (1989): −0.95

(Φf − Φ0 ) ∙ (e−p ∙ x − 1) ⎞ x ⎛ + RPOC = ⎜0.16∙ ⎜⎛a0 + ⎟ ω p∙ω∙ (Φf − 1) ⎝ ⎠ ⎝

⎞ ⎟ ∙POC ⎠

(2)

where a0 (kyr) is the apparent initial age of organic matter reflecting the reactivity of organic matter undergoing degradation in the seafloor. This adjustable parameter defines the distribution of POC degradation 3

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fPOC converts between POC (dry wt%) and DIC (mmol cm−3): fPOC = MWC·Φ/(10·(1-Φ)·ρS), where MWC is the molecular weight of carbon (12 g mol−1), ρS is the density of dry sediments and Φ is the porosity.

Fig. 2. Measured (symbols) and modeled (curves) profiles for 14C ages of TOC. Grey and black dots represent data from BC01 and GC02 respectively. The sediment burial velocity derived from modeling the 14C age (Eq. 1) is 8 cm kyr−1.

4. Results Sediments in upper 50 cm of GC02 consisted of brownish, homogeneous clay. From 50 to ~100 cm, the colour of sediments gradually changed from light brownish to light grey. Below ~100 cm to the end of the core (146 cm), the sediments were predominantly a grey to dark grey, foraminifera-rich ooze. The fitted porosity profile constrained by the measured data of BC01 is shown in Fig. S1. The sediment burial velocity estimated from 14C-AMS dating of TOC was 8 cm kyr−1 (Fig. 2 and Table 2). Down-core variations of TOC content, TOC/TN molar ratios, and δ13C and δ15N of POM are shown in Fig. 3. TOC generally increased with depth, ranging between 0.31 and 0.96% with an average of 0.62 ± 0.18 (n = 46). δ13C mirrored the TOC profile, decreasing from −21.3 to −23.2‰ with an average of −22.4 ± 0.51‰ (n = 46). Abrupt changes in TOC and δ13C from ~100 to 80 cmbsf were observed. Likewise, TOC/TN and δ15N varied from 8.2 to 12.9 and from 5.2 to 9.1‰, respectively, exhibiting marked gradient changes at ~80 cm. NO3− concentrations displayed near-seawater values in the upper 20 cm, and then decreased sharply and became depleted by ~90 cm (Fig. 4). NH4+ concentrations showed the opposite trend to NO3−, being depleted within the top 20 cm and enriched below it due to organic matter degradation. Likewise, DIC concentrations increased with sediment depth from 2.8 to 4 mM, whereas SO42− concentrations displayed only a slight decrease over the same depth.

Table 2 14 C-AMS dates of BC01 and GC02. Core

Depth (cm)

Age (calibrated, yr B.P.)

BC01

0–1 12–14 18–20 8–9 58–59 87–88 116–117 142–143

2067 ± 20 3596 ± 50 4491 ± 20 3433 ± 40 7471 ± 80 12,123 ± 190 16,556 ± 280 19,259 ± 390

GC02

a

RO2NH4 = kNH4+ ∙ [O2] ∙ [NH4+] Aerobic NH4 oxidation

NH4+ + 2O2 + 2HCO3− → NO3− + 3H2O + 2CO2

K NO− [SO2 −] KO2 3 ∙ ∙ 2− 4 /f [O2] + KO2 [NO− [SO4 ] + K SO2 − POC 3 ] + K NO− 3 4

(CH2O)(NH3)rNC + 1/2SO42− + rNCCO2 + rNCH2O → 1/2H2S + (1 + rNC)HCO3– + rNCNH4+ POM degradation via sulfate reduction

R deni = (CH2O)(NH3)rNC + 4/5NO3− → (1/5 – rNC)CO2 + 2/5 N2 + (4/5 + rNC)HCO3– + rNCNH4+ + (3/5 − rNC)H2O Denitrification

(CH2O)(NH3)rNC + O2 → (1 − rNC)CO2 + rNCHCO3− + rNCNH4+ + (1 − rNC)H2O POM degradation via aerobic oxidation

Rsr = RPOC ∙

[O2 ] /f [O2] + KO2 POC [NO− ] KO2 3 RPOC ∙ ∙ / fPOC [O2] + KO2 [NO− 3 ] + K NO− 3

Raer = RPOC ∙

Stoichiometry Process

Table 1 Biogeochemical processes and their kinetic rate expressions considered in the model.

Rate expressiona

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Fig. 3. Depth profiles of (A) TOC, (B) δ13C of sedimentary organic matter, (C) δ15N of sedimentary organic matter, (D) TOC/TN, and (E) estimated relative contribution of terrestrial organic matter (TOM) and marine organic matter (MOM) to the total organic fraction.

inorganic nitrogen on both TOC/TN and δ15N of bulk sediments. Plots of TN/TOC versus δ13C and δ15N versus δ13C thus enable the sources of organic matter to be elucidated and the contributions from marine and terrestrial organic matter to be assessed (e.g., Goni et al., 2008; Sanchez-Vidal et al., 2013). As shown in Fig. 5C and D, the POM in GC02 sediments is a mixture of terrestrial C3 organic matter and marine algae, but apparently with a greater contribution of marine organic matter. To quantitatively estimate the fractional contributions of terrestrial and marine organic matter to the bulk sedimentary POM pool, a two end-member mixing model based on δ13C was used:

5. Discussion 5.1. Organic matter provenance The composition of POM provides information on the mixing of the various organic matter sources and their subsequent degradation history. Differentiation of organic matter sources can be based on elemental and isotopic parameters provided that the influence of early diagenesis on the composition of POM has been minimal which is usually the case for δ13C and δ15N (Meyers, 1994; Middelburg and Nieuwenhuize, 1998). Organic matter in marine sediments is typically composed of a mixture of marine (autochthonous) and terrestrial (allochthonous) organic matter. In general, terrestrial vascular organic matter is more depleted in 13C than organic matter produced in marine surface water via photosynthesis (Emerson and Hedges, 1988; Hedges et al., 1997; Burdige, 2007). In addition, marine and terrestrial organic materials are characterized by distinct nitrogen isotopic compositions, whereby terrestrial organic matter is more depleted in 15N relative to marine algae (Fogel and Cifuentes, 1993; Meyers, 1997). Considering that effective methods of completely removing and extracting organic nitrogen from bulk sediments are lacking, δ15N of organic nitrogen is often equal to the δ15N of bulk nitrogen. This should not impede identification of the origins of organic matter and reconstructing past ocean chemistry and circulation if inorganic nitrogen only accounts for a small fraction of the bulk nitrogen (Ganeshram et al., 2000; Shigemitsu et al., 2008). The contribution of inorganic nitrogen to bulk nitrogen in GC02 is inferred to be minimal, as shown by a small TN intercept of the linear regression of the relationship between TOC and TN (Fig. 5A). Furthermore, inorganic nitrogen generally tends to lower the δ15N of bulk nitrogen (Freudenthal et al., 2001; Schubert and Calvert, 2001). The evident negative correlation between TOC/TN and δ15N suggests a decreasing contribution of terrestrial organic matter over time (Fig. 5B). Thus, the small TN intercept and negative correlation between TOC/TN and δ15N imply a negligible influence of

Fterr + Fmar = 1

(3)

δ13Cs = Fterr × δ13Cterr + Fmar × δ13Cmar

(4)

where Fterr and Fmar represent the fractions of terrestrial and marine organic matter, and δ13Cs, δ13Cterr, and δ13Cmar are carbon isotopic compositions of the sediment, and terrestrial and marine end-members, respectively. We used a value of −20‰ for the marine end-member and −28‰ for the terrestrial end-member according to previous data reported in the Gulf of Papua (Bird et al., 1995; Aller and Blair, 2004; Goni et al., 2008). Based on this mixing model, the POM in GC02 sediments was dominated by marine organic matter with a contribution increasing from 59% to 84% with decreasing depth (i.e. over the Holocene). The corollary of this finding is that the contribution from terrestrial organic matter decreases simultaneously from 41% to 16% (Fig. 3E). Because terrestrial organic matter is largely brought to the oceans by rivers, most is deposited and trapped on the shelf and only a small amount bypasses the continental margins to reach the deep ocean. Globally, it is estimated that the majority of the organic matter buried in marine sediments is of marine origin, especially in non-deltaic, continental margin and deep-sea settings (Gough et al., 1993; Burdige, 2005). Nevertheless, a number of studies have demonstrated that terrestrial organic matter input is somewhat underestimated in deep-sea 5

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Fig. 4. Measured (symbols) and modeled geochemical profiles (curves) for (A) POC, (B) O2, (C) NO3−, (D) NH4+, (E) DIC, and (F) SO42−. Black solid curves represent the steady-state model profiles using constant POC flux. Red dashed curves represent the non-steady-state model profiles using the POC flux calculated by Eq. (5). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

thus poised to accumulate a relatively large amount of terrestrial organic matter, possibly to abyssal and hadal depths (Luo et al., 2019).

settings, especially those connected to the margin by well-developed submarine canyon systems adjacent to river months (Holtvoeth et al., 2001; Puig et al., 2014). This is well exemplified by the Congo deep-sea fan system, through which a great amount of terrestrial organic matter is transferred from the Congo River to Angola Basin (Stetten et al., 2015; Baudin et al., 2017). In addition, mountainous rivers on highrelief islands along active margins coupled with narrow shelves and well-developed canyon systems have also been considered efficient in channeling terrestrial organic matter into the deep ocean (Lyons et al., 2002; Blair and Aller, 2012; Liu et al., 2016). The Markham River to the west of sampling site is a small mountainous river, but acts as the main conduit for transporting sediment from Papua New Guinea to the Solomon Sea owing to its steep- and short-reached catchment and high tropical rainfall (Milliman and Syvitski, 1992). The sediments can be transported SE-E downslope through its submarine continuation via the Markham Canyon, and may even reach the western end of the New Britain Trench in gravity flows and nepheloid layers (Davies et al., 1986; Whitmore et al., 1999). Other smaller canyons that are fed by the rivers on the central New Britain Island, e.g., Andru Canyon and Johanna Canyon, are likely to supply terrestrial materials to the sampling site from the north (Whitmore et al., 1999). It is therefore conceivable that up to 40% of the organic matter in GC02 originates from the land in spite of the relatively large water depth (~4000 m). Similarly, terrestrial material has been reported in the bathyal and abyssal oozes of the Woodlark Basin, southwestern Solomon Sea (Alongi, 1990 & Alongi, 1992). The Solomon Sea is surrounded by active margins and large high-relief islands under the influence of warm, humid climate within the domain of WPWP, and

5.2. A decrease in POC deposition at the deglaciation period/Holocene transition Variations in sediment TOC contents are usually ascribed to temporal changes of sedimentation rate and/or POC flux. Based on the 14CAMS dating results of TOC in BC01 and GC02, sediment deposition appeared to be relatively stable for the last ~20 ka with an average sedimentation rate of ~8 cm kyr−1. Therefore, we postulate that the variation in TOC contents was mainly caused by the fluctuations in POC rain rate. Considering the rather low TOC content in the sediment, a change of POC flux without an attendant change in sedimentation rate is plausible. With regard to examining the change of POC rain rate to the seafloor, we first assumed that POC approximates to TOC, considering that the dissolved organic carbon accounts for a negligible fraction of TOC in the sediments. One conspicuous feature of the POC profile is the sharp decrease of POC content from 0.82 to 0.52% between ~100 and 80 cmbsf (Fig. 3A). This is indicative of a non-steady-state deposition of POC and usually interpreted as a transient change in POC rain rate (e.g., Mogollón et al., 2012; Wehrmann et al., 2013; Luo et al., 2015). According to the sediment burial velocity derived from 14C dating of TOC, there may have been a decrease in POC flux to the seafloor roughly from the deglaciation period to early Holocene. Based on these observations, we used an inverse non-steady-state model to reconstruct the flux fluctuation of POC to the seafloor in the WPWP. We began by 6

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Fig. 5. Cross plots of (A) TN vs. TOC, (B) TOC/TN vs. δ15N, (C) TN/TOC vs. δ13C, and (D) δ15N vs. δ13C of sedimentary organic matter. MOM and C3 represent compositional ranges of marine organic matter and C3 vascular plants, respectively (Meyers, 1994; Middelburg and Nieuwenhuize, 1998; Goni et al., 2008). The arrows indicate a trend towards a greater contribution of terrestrial C3 vascular plants to the bulk sedimentary organic matter with depth in the sediment. Table 3 Depth-integrated rates from the model.

Total POC mineralization rate POC degradation via aerobic oxidation Denitrification POC degradation via sulfate reduction Aerobic NH4 oxidation

GC02 (prior to 13 ka)

GC02 (present day)

Unit

1.76 0.34 0.17 0.58 0.16

1.22 0.21 0.12 0.42 0.91

μmol cm−2 yr−1 of C μmol cm−2 yr−1 of O2 μmol cm−2 yr−1 of N μmol cm−2 yr−1 of S μmol cm−2 yr−1 of N

determining the POC flux and initial age of POC (a0) in the pre-Holocene setting. To achieve this, the model was run to steady-state until the modeled profile simulated the POC data below ~100 cmbsf (Fig. 4). This required a POC flux of 75 μg cm−2 yr−1 and an initial age of organic matter of 10 kyr (Table S2). For the pre-Holocene setting, the total POC mineralization rate was 1.76 μmol cm−2 yr−1 (Table 3). However, porewater profiles of NO3−, NH4+, and DIC were not well fitted in this scenario due to the relatively high organic matter content and mineralization rate (Fig. 4). Subsequently, we attempted to simulate the measured POC profile by decreasing the POC flux (FPOC) from the deglaciation period (~13 kyr B·P.) over several thousand years using the following function:

FPOC (t ) = a1 − a2 × (Erf[(t − b1)/ b2])

initial conditions, the transient decrease in POC contents was tuned by adjusting parameters a1, a2, b1, and b2 (ensuring that a1 + a2 = 75 μg cm−2 yr−1) and was further constrained by porewater profiles of NO3−, NH4+, and DIC. The results show that both POC and porewater data are consistent with a decrease in POC rain rate from 75 to 37.5 μg cm−2 yr−1 over 4 kyr followed by a constant POC flux for the last 9 kyr (Figs. 4 & 6). Note that although the change of POC flux persisted for 4 kyr, the sharp decrease in POC flux occurred from ~12 to ~9.5 ka. Under this present-day condition, the depth-integrated rate of total POC mineralization (1.22 μmol cm−2 yr−1) is lower compared to the pre-Holocene setting (Table 3). This hindcast transient modeling thus allows the POC rain rate and its degradation from the deglaciation period to early Holocene to be quantitatively reconstructed. Further tests with the model (not shown) revealed that it was not possible to simulate the decrease in POC contents by increasing instead the reactivity of organic matter. In this case, elevated organic matter mineralization after 13 ka leads to greater accumulation of metabolites in the porewater and a clear offset between the simulated and measured

(5)

where a1 + a2 represents the pre-Holocene POC flux, 2 × a2 represents the decrease in POC flux, a1 − a2 represents the present-day POC flux, b1 is half of the perturbation length and b2 controls the rate of change of the POC flux. Using the steady-state model results as the 7

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of organic matter as well as a slight decrease in TOC/TN (Fig. 3). This suggests that although the content of bulk sedimentary organic carbon declined, the relative contribution of marine organic matter to the bulk sedimentary organic matter increased during the transition from the deglaciation period to early Holocene. This disagrees with the conclusions of Burke et al. (1993) and Barash and Kuptsov (1997) that a decline in primary productivity prevailed from the end of Pleistocene to the early Holocene. In fact, we speculate that the contribution of terrestrial organic matter input has diminished since the deglaciation period (~13 ka) (Fig. 3E). It is likely that a rapid rise in sea level from the deglaciation period to early Holocene reduced the supply of terrigenous material. Given its proximity to both New Britain and Papua New Guinea narrow and steep margins, a well-develop submarine canyon system, and high precipitation, a sea level fluctuation of several tens of meters would probably be the dominant factor governing the delivery of terrestrial organic material to the sampling site. A conspicuous decline in POC accumulation rates during the deglaciation period and early Holocene has also been observed in the mid-slope of the northeastern Pandora Trough in the Gulf of Papua and in the northern Gulf of Mexico. This change was primarily attributed to a reduced delivery of terrigenous organic matter caused by post-glacial sea level rise (Febo et al., 2008; Ingram et al., 2013). In contrast, intensified upwelling in the northwestern South China Sea and increased freshwater input into the Kara and Laptev Seas continental margin (eastern Arctic Ocean) at the deglaciation/Holocene transition led to enhanced surface water productivity and increased sedimentary POC content (Boucsein et al., 2002; Luo et al., 2015). Consequently, the trend in the evolution of sedimentary POC may vary from region to region. Based on our results, we postulate that the marked decrease in sedimentary TOC content for a period of around 4 kyr since the deglaciation period in the southern WPWP was mainly caused by a reduced supply of terrestrial organic matter in response to a rapid rise in sea level. The decline in primary productivity presumably played a minor role in inducing the reduced POC flux and POC content at the seafloor.

Fig. 6. Change of POC flux at the seafloor during the last 13 kyr using Eq. (5). POC flux decreased from 75 to 37.5 μg cm−2 yr−1 over ~4 kyr followed by a lower constant POC flux for the last ~9 kyr.

DIC and NH4+ profiles. This is evident in the pre-Holocene scenario shown in Fig. 4 where elevated concentrations of DIC and NH4+ result from increased rate of POC degradation due to too high POC rain rate. Therefore, the decline in POC flux to the seafloor is the most likely and plausible explanation for the decrease in TOC contents at 80–100 cm depth. 5.3. Possible mechanisms for the change of POC flux during the deglaciation period and early Holocene Primary productivity in the WPWP decreased strongly from the west to east and showed prominent glacial-interglacial cyclicity in the western sector. The central and eastern sectors, however, exhibited little difference in the primary production during the late Pleistocene (Kawahata, 1999). This was mainly due to deeper thermocline in the central and eastern sectors of WPWP that prevented upwelling of nutrient-rich deep water in spite of intensified wind-induced mixing and upwelling during the glacial periods (Kawahata, 1999). It is thus likely that the production, degradation, and burial of organic matter in the WPWP vary in response to the localized geography despite being within a similar oceanographic regime. This observed sharp increase in POC at the Holocene boundary is consistent with the TOC content in sediment cores collected from the southeastern Solomon Sea and the Bismarck Sea. This has been hypothesized to reflect a regional feature and might represent a paleoceanographic event in the equatorial western Pacific (Barash and Kuptsov, 1997). According to our modeling efforts, this increase in POC content could be reconstructed by increasing POC flux to the seafloor from 37.5 to 75 μg cm−2 yr−1 at the boundary of Holocene and deglaciation period. In fact, Burke et al. (1993) speculated that a significant drop in the overall abundance of benthic foraminifera in deep-sea sediments of Ontong Java Plateau (western WPWP) was caused by a marked decrease in the supply of organic matter during the deglaciation period. This apparently regional decrease in the flux of organic carbon at the transition from the deglaciation to early Holocene has been attributed to a decline in surface productivity (Burke et al., 1993; Barash and Kuptsov, 1997). Indeed, surface productivity and export production has a close linkage with the deposition and burial of organic matter in the sediments (Suess, 1980). The flux of POC to the seafloor is routinely taken as a useful indicator for surface productivity under the assumption that a majority of sinking organic material is produced in the surface ocean. Nevertheless, terrestrial organic carbon can account for a non-negligible component of sedimentary organic carbon in some continental marginal and even deep-sea settings (e.g., Goni et al., 1997; Stetten et al., 2015; Luo et al., 2017 & Luo et al., 2019). In GC02, the obvious decline in TOC content is accompanied by an increase in δ13C and δ15N

6. Conclusions We reconstructed a sharp decrease in POC content in the deep-sea sediments of the northwestern Solomon Sea combining measured POC and porewater data and a non-steady-state reaction-transport modeling. Hindcast model results suggest that POC content and porewater profiles (NO3−, NH4+, DIC and SO42−) can be simulated with a reduction in POC rain rate from 75 to 37.5 μg cm−2 yr−1 at the onset of the deglaciation period (~13 ka) for a duration of ~4 kyr. We attributed this decrease to reduced terrestrial organic matter input driven by the rapid rise in sea level rather than to reduced primary productivity or a change in the reactivity of bulk organic matter. A decrease in the relative contribution of terrestrial organic matter input at this time was inferred based on the analysis of carbon and nitrogen contents and isotopic compositions of the sedimentary organic matter. Consequently, we highlight the important role of high-relief mountainous islands in the WPWP in supplying terrestrial organic material to the deep ocean, especially during sea level lowstands. Acknowledgments We thank the captain and crews of RV Zhangjian for their great help with sampling at sea. We express our gratitude to Michael E. Böttcher for handling the manuscript and two anonymous reviewers for their constructive commentaries. This study was financially supported by National Natural Science Foundation of China (Grant No. 41703077), Qingdao National Laboratory for Marine Science and Technology (Grant No. QNLM2016ORP0208), the Shanghai Sailing Program (Grant No. 17YF1407800), and the Strategic Priority Research Program of the Chinese Academy of Sciences (Grant No. XDB06030102). 8

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