Carbon burial records during the last ~40,000 years in sediments of the Liaohe Delta wetland, China

Estuarine, Coastal and Shelf Science 226 (2019) 106291

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Carbon burial records during the last ~40,000 years in sediments of the Liaohe Delta wetland, China

T

Guangming Zhaoa,b,c,d, Siyuan Yea,b,d,*, Edward A. Lawse, Lei Hea,d, Hongming Yuana,d, Xigui Dinga,d, Jin Wanga,d a

The Key Laboratory of Coastal Wetlands Biogeosciences, China Geologic Survey, Qingdao, 266071, PR China Laboratory for Marine Geology, Qingdao National Laboratory for Marine Science and Technology, Qingdao, 266061, PR China c Shandong University of Science and Technology, Qingdao, 266590, PR China d Qingdao Institute of Marine Geology, China Geologic Survey, Qingdao, 266071, PR China e School of the Coast & Environment, Department of Environmental Sciences, Louisiana State University, Baton Rouge, LA, USA b

ARTICLE INFO

ABSTRACT

Keywords: Liaohe delta Wetlands Sedimentary processes Carbon sequestration CIA Late pleistocene

Delta wetland sediments constitute a long-term natural carbon sink and play a critical role in the global carbon cycle. In this study, borehole core ZK3 (36.7-m long), drilled in 2012 in the Liaohe Delta wetland, was investigated to assess the rate of carbon sequestration and the factors influencing carbon burial since the Late Pleistocene. Here we report the results of integrated analyses of the core, including its sedimentary lithology, grain size, foraminiferal abundance, chemical elements, and accelerator mass spectrometry (AMS)14C and optically stimulated luminescence (OSL) dates. The sedimentary environment has evolved from a fluvial-deposit, limnetic-deposit, littoral-deposit, shallow sea–deposit, and finally to a delta-deposit environment since 40,000 cal yr BP. Environmentally induced differences in apparent mass accumulation rates (AMARs) of organic carbon (OC) have been significant; they have ranged between 3.73 and 30.77 g/(m2·yr). The fact that the highest rates were associated with the delta-deposit environment indicates that the rate of carbon sequestration was greater in the sediments of estuarine wetlands than sediments of the continental shelf. The chemical index of alteration (CIA) proxy responded to several cold events, and there was a positive correlation between the CIA and OC-AMAR. Climate may therefore have regulated the excursions of ecosystem productions and in turn impacted the dynamics of carbon burial.

1. Introduction With the increasing global average atmospheric carbon dioxide (CO2) concentration, a single emission–reduction strategy has been switched to a plan that combines reducing anthropogenic emissions of CO2 (mitigation) with strengthening carbon (C) storage of natural ecosystems because of their high C sequestration rates and capacity for carbon storage (Canadell and Raupach, 2008). As an important part of blue carbon sinks, coastal wetlands play a critical role in carbon sequestration and mitigation of global warming effects (Bridgham et al., 2006b; Chmura et al., 2003; DeLaune and White, 2012; Mcleod et al., 2011). The organic carbon pool in sediments may exceed the amount of carbon in living vegetation by a factor of 2–3 (Lettens et al., 2005; Schlesinger, 1990). Sediment carbon sequestration is believed to be one of the cost-effective ways to mitigate the adverse effects of atmospheric CO2 increases, including global warming (Naizheng et al., 2011; Olson,

*

2013; Sommer and Bossio, 2014). Costal wetland sediments constitute a long-term natural carbon sink and thus affect atmospheric CO2 concentrations (Kroeger et al., 2017). High rates of C accumulation in coastal sediments are associated with high plant productivity and burial of organic matter through sedimentation. Substantial spatial variability of wetland C accumulation rates has been related to differences in tidal range, elevation, freshwater input, and sediment availability (Craft and Richardson, 1998; Kirwan and Guntenspergen, 2010; Unger et al., 2016). The two species of carbon in the environment are organic carbon (OC) and inorganic carbon (IC) (Koziorowska et al., 2017, 2018). The material flux reaching costal wetland sediments comprises a mixture of these two forms. Whether they are autochthonous or allochthonous, both OC and IC are readily deposited in bottom sediments (Teske et al., 2011). The former are intensively mineralized in the water column and at the sediment—water interface (Holding et al., 2017; Teske et al., 2011),

Corresponding author. The Key Laboratory of Coastal Wetlands Biogeosciences, China Geologic Survey, Qingdao, 266071, PR China. E-mail addresses: [email protected], [email protected] (S. Ye).

https://doi.org/10.1016/j.ecss.2019.106291 Received 31 July 2018; Received in revised form 11 July 2019; Accepted 11 July 2019 Available online 12 July 2019 0272-7714/ © 2019 Published by Elsevier Ltd.

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Fig. 1. Schematic map of the Liaohe Delta and the location of core Zk3

where the redox conditions determine the efficiency of mineralization and biogenic element exchange between sediments and the overlying water (Ingall et al., 2005; Jørgensen et al., 2005). Deltaic coastal wetlands experience vegetation and habitat shifts associated with changes in relative sea level, river flooding, channel migration, and storm deposition or erosion. Therefore, rates of C accumulation will change over time concomitant with changes in sedimentation, hydrodynamics, freshwater availability, and productivity. Within the sediment profile, loss of OC can occur through microbial respiration and erosion and tidal export, whereas storm deposition and inputs of fresh organic matter will lead to increases (DeLaune et al., 2018). The burial of organic matter during the postglacial period is linked with paleoenvironmental changes (Yang et al., 2011). Quantitative analyses of carbon burial records have involved the determination of apparent mass accumulation rates, which have helped to clarify stratigraphic divisions, processes associated with evolution of the paleoenvironment, and the associated timespans. Quantification of C accumulation rates may depend on sediment compaction over time and consequently require a bulk density correction to validate comparisons between depths and among different systems. Sediment accumulation rate, which is inversely correlated with % OC, is a major factor

controlling the organic carbon flux to sediments (Cui et al., 2016; Ramirez et al., 2016; Zhao et al., 2015) (Cui et al., 2016; Ramirez et al., 2016; Zhao et al., 2015). The goal of this study was to provide quantitative information about long-term C burial in the Liaohe Delta (LHD), the northernmost coastal wetland in China. Combined with core elevation and sea level change data, sedimentary lithology, grain size, radiocarbon dating and foraminiferal abundance were analyzed to establish the chronostratiographic framework of the LHD since ~40,000 year BP of the late Pleistocene. Carbon burial and the factors that influence carbon burial are discussed in the context of this information. 2. General setting The study site is located in the LHD (121°25′E–122°55′E, 40°40′N–41°25′N) in northeastern China. The LHD has been formed by sediments deposited by the Liaohe River, Daliao River, Daling River, and Xiaoling River (Fig. 1). Of these rivers, the Liaohe River is the largest, with a total length of 1396 km and a drainage area of 2.19 × 105 km2. The Liaohe River was previously named the Shuangtaizi River, but the name was changed to Liaohe River in early March of 2

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2011. The Daliao River is the largest contributor to runoff, whereas the Daling River is the largest contributor to sediment discharge (He et al., 2018). The Delta has a total area of about 5200 km2, of which about 1280 km2 was historically covered by wetlands (Li et al., 2012). About 786 km2 of the wetland is marsh vegetated by common reed (Phragmites australis (Cav.) Trin.EX Steud). The reed marsh in the Liaohe Delta represents probably the largest reed field in the world (Brix et al., 2014). It has been designated the Shuangtaizihekou National Nature Reserve since 1986 and has also been listed as a Ramsar Site since 2004 (Li et al., 2012). The Liaohe River and Daliao River have not been isolated throughout their history. According to historical records, the Liaohe River and Daliao River entered the sea through the same channel from the Han Dynasty (206 BCE to 24 A.D.) to 1861 CE (Pan, 2005). They partially separated after a crevasse formed in the right bank of the Liaohe River in 1861. The crevasse led to part of the Liaohe River's entering Liaodong Bay by Panjin, while the rest of the river still entered the bay by Yingkou. The eastern portion of the Liaohe-Daliao River catchment is mainly mountainous, with high vegetative coverage and relatively high precipitation. This eastern area yields about 70% of the annual runoff from the combined catchments (Wang, 2012). In contrast, the vegetative coverage of the western area is < 30%; the result is severe soil loss, and the western area is therefore the main source to sediments in the catchment (Zhao et al., 2008). About 83% of the sediment in the Daling River is derived from the upstream portion of its catchment (Li et al., 2000). The sediments in the Xiaoling River are also derived mainly from its upstream catchment, as are about 85% of the sediments in the Xiaoling River (Zhang et al., 1995). The estuary of the four rivers is located in Liaodong Bay, Bohai Sea, which is a semi-enclosed epeiric sea (Fig. 1). The waves in Liaodong Bay are primarily driven by the wind. The average height of the waves is less than 1 m, with 5 m as the maximum height (Liu, 1996). The average tidal range is about 2.7 m at the Liaohe River mouth (Zhu et al., 2007). Tidal water can reach up the Liaohe River to Panjin, which is about 61 km from the river mouth. The alongshore currents of the northernmost part of Liaodong Bay are much weaker than the tidal currents. Except for some individual months (July or August), the coastal currents move in a clockwise direction (Zhao et al., 1995). The climate of the study region belongs to a warm temperate continental monsoon climate. The annual average temperature is 8.3–8.4 °C. The annual average rainfall is 612–640 mm, and evaporation is 139–171 cm (Yang et al., 2008).

(AMS) and optically stimulated luminescence (OSL); and identification of foraminifera. Subsamples with a thickness of 10 cm were taken, generally at 20-cm intervals, for analyses of TC, OC, and other parameters. A black, rigid, 5 cm–diameter plastic cylinder was pushed into the core for OSL sampling. The sample was sealed and stored in a shockproof container. The 14C ages of six samples were measured by Beta Analytic Inc. (Miami, Florida, USA) using accelerator mass spectrometry (AMS). The radiocarbon ages were corrected for the regional marine reservoir effect (ΔR), which was considered to be −178 ± 50 yr for the study sea area (Southon et al., 2002), and calibrated using Calib Rev. 5.0.1 (Stuiver et al., 2005). Table 1 gives calibrated ages ± one standard deviation (1σ) uncertainty. Calibrated ages in this paper are reported as years before AD 1950 (cal yr BP), and uncalibrated ages are given as 14C years BP (14C yr BP). The average interval of microfossil sampling was ~0.25 m in the cores, with a smaller sample interval (~0.10 m) in some layers. Samples were picked and studied under a Zeiss optical stereoscope, using the method described by He et al. (2018). The “foraminiferal abundance” parameter in this paper is the number of foraminifera per 50 g of dry sediment. The “simple diversity” is the number of foraminiferal species in each 50 g-sample. One sample was collected from at the depth of 23.5 m for OSL dating at the Testing Center of Qingdao Institute of Marine Geology using a Daybreak 2200 TL/OSL reader (Table 1). The sample exhibited high OSL sensitivity and a well-behaved OSL signal. The dating was performed using the method of Wintle (1997) and Qiu et al. (2014). An age-elevation model was established to facilitate linear interpolation between measured ages. The model assumed that the sedimentation rate was constant between dated levels. Thus, the ages estimated on the basis of the age-elevation model for a given strata where dating information was absent are referred to as model ages in this manuscript (Liu et al., 2017). Samples were oven dried at 60 °C, passed through a 2-mm sieve after being ground, and analyzed using a non-aqueous volumetric method for TC content. The concentrations of OC were determined in a similar way after treating another set of subsamples with 2% diluted phosphoric acid before analysis to remove inorganic carbon. Concentrations of IC were equated to the difference between TC and OC. Before major elemental analysis, about 0.05 g of each sample was quantified, and 1.5 mL of HNO3 and 1.5 mL of HF were used to dissolve the sample at 190–200 °C for 48 h. Then 50 g of the liquor was prepared for measurement. The contents of Al2O3, CaO, Na2O, and K2O were measured using an Inductively Coupled Plasma Optical Emission Spectrometer (ICP-OES). Duplicate measurements of the geostandard GSM09 showed that the percent differences between measured and certified values were generally less than 5% in magnitude. Grain sizes of the sediments were measured at the Experimental & Testing Center of the Qingdao Institute of Marine Geology, China Geological Survey, Qingdao, China. An appropriate amount of sample was placed into a beaker along with 15 ml of H2O2 solution (30%). The

3. Samples and methods Borehole core ZK3 (40° 52′4.59″N, 121° 36′2.07″E) (Fig. 1) was obtained in the LHD in May 2012 by rotary drilling. The core was 36.7 m long; the top was 2.73 m above sea level, and the average recovery was 91.5%. A GJ-240-type drilling rig and boring casing were used for the coring. In the laboratory, the core was split in half, described, and subsampled for analyses of bulk density, chemical components, and grain size; dating via accelerator mass spectrometry Table 1 AMS 14 C dating and optically stimulated luminescence dating of Core ZK3. Sample number

ZK3S1 ZK3S9 ZK3S25 ZK3S3 ZK3S27 OSL-1 ZK3S8

Depth (m)

3.5 11.2 11.95 12.83 16.05 23.75 29.05

Elevation (m)

−0.77 −8.47 −9.22 −10.10 −13.32 −21.02 −26.32

Method

14

AMS C AMS14C AMS14C AMS14C AMS14C OSL AMS14C

Materials

Mollusk shell Mollusk shell Mollusk shell Plant fragment Mollusk shell

δ13C (‰)

Species

Mactra veneriformis Corbicula fluminea Ruditapes philippinarum Biomphalaria sp.

Plant fragment

3

Conventional

14

0 −9.6 −0.4 −25.3 −12.2

390 ± 30 7640 ± 30 4390 ± 30 9980 ± 40 12850 ± 50

−22.7

35990 ± 340

C age (aBP)

Calibrated age (cal. a BP) Intercept

Range (1σ)

190 8425 4775 11425 15320 24500 40625

130–274 8400–8446 4693–4845 11308–11410 15206–15414 22000–27000 40246–41030

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sample was soaked in the H2O2 for 24 h to remove organic matter. The sample was then bathed in 5 ml of HCl solution (3 mol/L) for 24 h to remove calcareous cement and shell materials. Grain size was determined using a laser particle size analyzer (Mastersizer 3000, Malven Instruments Ltd., UK) that could distinguish particles in the size range 0.02–2000 μm; the resolution ratio was 0.01 Ф, and the analytical error was better than ± 2%. A subsample of approximately 700 g from each of the sediment samples was weighed to the nearest 35 g on a Shuangjie electronic balance (model JJ3000) and then oven dried at 105 °C for a minimum of 24 h to a constant weight. Moisture content was equated to the loss in weight, and the weight after drying (dry weight) was used to calculate the bulk density (BD) of the samples. The apparent mass accumulation rates of the OC, TC, or IC were calculated from the product of OC, TC, or IC content, bulk density, and the apparent sedimentation rate (ASR) of a sample as follows (Liu et al., 2017): Ci-AMAR = Ci × ASR × BD × 10

4.1.1. Fluvial deposit (U5) (depth in core: 23.4–36.7 m; elevation: −20.67 m to −33.67 m) This unit was composed mainly of silt and fine-grained sand, with mean grain size Ф ranging between 2.2 and 6.56. The lower part is dominated by greyish clayey silt intercalated grey silt and fine-grained sand, whereas the upper part consists of grey middle fine sand with boulder clay and calcareous concretions. The absence of foraminifera in this unit indicates that it is a fluvial deposit. The much coarser grain size in the upper section than in the lower section suggests that the upper section is mostly a channel fill deposit and the lower is a flood plain deposit. The AMS 14C age from plant fragments in the middle of U5 (core depth: 29.5 m) was 40,625 cal yr BP, which is beyond the effective measurement range of AMS14C and relatively old in this layer. We therefore discarded this age. The OSL age (24.5 ka) at a depth of 23.75 m was reliable and reasonable. 4.1.2. Limnetic deposit (U4) (depth in core: 12.4–23.4 m; elevation: −9.67 to −20.67 m) This unit was dominated by grey to greyish clayey silt, intercalated with several silt to silty sand beds. The transitions between U4 and upper and lower strata were abrupt. The absence of benthic foraminifera suggests a terrestrial environment. The relatively uniform lithology of U4 compared to U5 was indicative of a relatively stable environment. The obvious variation of the mean grain size Ф, 3.95–7.1, may have resulted from back-and-forth shifts of the main river channel relative to the core position. The fine grain sizes in this deposit may be attributable to weak hydrodynamic conditions after the channel shifted away from the core site. The interpretation is that U4 was deposited in a limnetic environment. Plant fragments at a depth of 12.83 m and mollusk shells at 16.05 m were dated to 11,425 cal yr BP and 15,320 cal yr BP, respectively.

(1)

where, Ci-AMAR is the OC, TC, or IC apparent mass accumulation rate (g/[m2·a]); Ci is the OC, TC, or IC content (mg/g); ASR is the apparent sedimentation rate (cm/a); and BD is bulk density of the sediment (g/ cm3). 4. Results and discussion 4.1. Depositional units, sedimentary environments, and dating of core ZK3 On the basis of the characteristics of the lithofacies and downcore distributions of benthic foraminiferal assemblages, the sediments in core ZK3 could be divided into five depositional units, designated U1–5 in descending order (Fig. 2).

4.1.3. Littoral deposit (U3) (depth in core: 11.8–12.4 m; elevation: −9.07 to −9.67 m) This unit was characterized by grey, silty sand and the occasional

Fig. 2. Comprehensive geological log of core ZK3. 4

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presence of small oysters. The transition between U3 and U4 was clearly erosional, and there were several erosional surfaces between the sandy beds of the lower section of U3. Mean grain size Ф ranged between 3.82 and 4.77. The abundance and simple diversity values of the benthic foraminifera reached maximum values of 160 and 16, respectively, and displayed a tendency to increase toward the surface. The benthic foraminifera assemblages were dominated by Ammonia beccarii vars., a species characteristic of low-salinity environments, and Protelphidium tuberculatum. This unit was interpreted as a littoral deposit at the beginning of the Holocene transgression. A mollusk shell at a core depth of 11.95 m was dated to 4775 cal yr BP. Combined with the dating result of 11,425 cal yr BP at a depth of 12.83 m in U4, the depositional time of the lower boundary of U3 was deduced to be 8237 cal yr BP by interpolation, which is close to the start time of the third transgression (8910 cal yr BP) in the Yellow River Delta (Saito et al., 2001). U4 therefore belonged to the postglacial transgression deposition of Marine Isotope Stage 1.

occur in a shallow-sea environment but instead reflected delta deposition. U14 was therefore interpreted as a prodelta deposit. 4.1.7. Delta front (U13) (depth in core: 4.7–10.2 m; elevation: −1.97 to −7.47 m) U13 was further subdivided on the basis of grain size into two sections, an upper section (4.7–7.3 m) and a lower section (7.3–10.2 m). The upper section was composed mainly of grey to dark grey fine sand with a high moisture content and black organic interbeds, and the lower section alternated between grey clayey silt and greyish silty fine sand with lenticular beddings. Some shell fragments were found at a depth of 6.5 m. The abundance of benthic foraminifera averaged 33 (range: 1–103), and the simple diversity was 16 (range: 4–22). This unit contained the maximum percent sand (99.91%) and the minimum percent clay (0%) in the whole core. Mean grain size Ф averaged 5.15. The coarser grain size of U13 relative to the underlying prodelta (U14) suggested that U13 was a delta front deposit according to the superposition relationship of a subaqueous delta. This unit was also subdivided into a mouth bar and distal bar because the sediments in the upper and lower parts were coarse and fine, respectively.

4.1.4. Shallow sea deposit (U2) (depth in core: 11.0–11.8 m; elevation: −8.27 to −9.07 m) This unit was composed of greyish fine sand interlaminated with clay layers and lenticular bedding. The transitions between the upper and lower strata were gradual. Mean grain size Ф values were 5.03–6.07. The abundances (111–240 with an average of 170) and simple diversities (21–22 with an average of 21.5) of the benthic foraminifera of this unit were the highest among all depositional units of the core. The benthic foraminifera assemblages were dominated by Elphidium magellanicum, Ammonia beccarii vars. and Protelphidium tuberculatum. Elphidium magellanicum is typically found in relatively lowtemperature and low-salinity habitats (Wang and Cheng, 1988) and is an indicator of nearshore, shallow-sea environments. It is distributed in the Bohai Sea mainly where the depth of the water is less than 20 m (Chen et al., 2008). The greater diversity and abundance of benthic foraminifera in U2 compared with the other units (Fig. 2) indicated that it was associated with a shallow-sea environment with stable hydrodynamic conditions and low ASRs. A mollusk shell at a core depth of 11.2 m was dated to 8425 cal yr BP, which was inconsistent with the date (4775 cal yr BP) of the underlying layer of U3 at 11.95 m. The explanation may be that wave action transported shells from the underlying, older strata to the upper layer. We accepted 4775 cal yr BP at a depth of 11.95 m and abandoned 8425 cal yr BP at a depth of 11.2 m.

4.1.8. Lower delta plain (U12) (depth in core: 2.45–4.7 m; elevation: 0.28 to −1.97 m) This unit was characterized by greyey clayey silt, intercalated with grey silt and silty fine sand layers. The mean grain size Ф averaged 5.1, and the grains became finer toward the surface. The moisture content of U12 was high, and its basal contact with the underlying unit was erosional. A single-lobe shell 1.5 cm × 2 cm in size appeared at a depth of 3.5 m, and some shell fragments (5–8 mm in size) were found at a depth of 3.74–3.95 m. Tidal bedding (i.e., wavy and lenticular bedding) was very common, and plant material and organic matter enrichment were observed at some depths. The abundance and simple diversity averaged 16 and 8.4, respectively. The assemblages were dominated by Ammonia beccarii vars. Unit U12 was probably deposited in a lower delta plain environment in one part of a delta plain. The sand beds of the lower section with tabular cross-bedding reveal a tidal channel deposit. The near absence of bioturbation and paucity of burrows in this unit may reflect high-energy conditions and high sediment accumulation rates (e.g., Abrahim et al., 2008) in the depositional environment. This scenario is consistent with the relatively low abundance and simple diversity of benthic foraminifera. 4.1.9. Upper delta plain (U11) (depth in core: 0–2.45 m; elevation: 2.73 to 0.28 m) The top interval (depths: 0–1.33 m) consisted of artificial fill. The rest of this unit was composed mainly of brown yellowish clayey silt interbedded with grey silt laminations, including decayed plant roots. No bioturbation was evident. Mean grain size Ф averaged 5.92. The average abundance of benthic foraminifera, 1.3, was low; single species were common. U11 was interpreted as an upper delta plain, also one part of a delta plain. The occurrence of a small number of benthic foraminifers in this unit may be due to seawater's reaching here over the high tide line during storm surge. Core ZK3 preserved a complete transgression–regression depositional sequence since the late Pleistocene in the following order: fluvial and limnetic deposit, littoral and shallow sea deposit, delta plain. Before 24,500 cal yr BP, fluvial deposits were distributed in the study area, and then the deposits gradually changed to limnetic deposits. Both the fluvial and limnetic deposits were terrestrial deposits corresponding to the low stand system tract (LST). The Holocene transgression reached the region at ~8237 cal yr BP, and a littoral environment started to develop widely. With the further rise of sea level, the study area gradually changed from a littoral environment to a shallow-sea environment for a long period of 3971 years. Both littoral and shallow-sea deposits were influenced by the Holocene transgression corresponding to the transgressive system tract (TST). When sea level stabilized, a

4.1.5. Delta deposit (U1) (depth in core: 0–11 m; elevation: −2.73 to −8.27 m) This unit was characterized by (1) clayey silt interbedded with silt or silty fine sand layers, (2) silt to fine sand, (3) fine sand, and (4) clayey silt with silty lenses from bottom to top. The pattern of grain sizes: fine →coarse→fine from bottom to top, was indicative of a deltaic sequence. Moreover, the abundance of benthic foraminifera and the content of shallow-sea benthic foraminifera species decreased while euryhaline species increased toward the top of this unit suggested a gradual transition toward a less saline habitat. The characteristics of U1 were therefore generally within the range typical of delta deposits. On the basis of its lithofacies characteristics and distribution of grain sizes, U1 could also be subdivided into four sections. 4.1.6. Prodelta deposit (U14) (depth in core: 10.2–11 m; elevation: −7.47 to −8.27 m) U14 consisted of grey clayey silt with greyish silty lenses. The average content of clay (7.79%) of this unit was the highest and sand (23.33%) was the lowest among all the depositional units of the core. Mean grain size Ф averaged 6.56. The abundance of benthic foraminifera averaged 61 (range: 13–109). The simple diversity was 16.5 (range: 12–21). The abrupt decrease of both the diversity and abundance of benthic foraminifera compared to the underlying unit indicated a considerably higher ASR in U14. This high ASR surely did not 5

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Fig. 3. Vertical distribution of OC, TC, IC, BD, OC-AMAR, TC-AMAR and IC-AMAR in core ZK3.

and the BD of the prodelta (U14) was only 1.76 g/cm3. Bulk density decreases with increasing organic matter content, and soils with a low bulk density appear to support better vegetative growth (Nair et al., 2001). The abundant small pores and great surface energy of fine-grained sediments can aid in absorbing and preserving organic carbon (Zhang et al., 2009). As shown in Fig. 4, the contents of OC, IC, and TC were significantly correlated with the clay content of the sediment (Fig. 4; P < 0.01). The higher C concentrations in the finer sediments compared to the coarser sediments suggest that high percentages of fine grain sizes favor retention of C. A similar observation has been made by previous researchers in the sediments of estuaries and of coastal regions along both the east and west coast of India (Anbuselvan et al., 2018; Godson et al., 2018; Gopal et al., 2017; Kasilingam et al., 2016; Nethaji et al., 2017). Sediment grain size is one factor that controls the abundance of C in sediments. Nevertheless, the contents of OC and IC in sediments are influenced by other factors associated with the complicated sedimentary environment in the LHD, which remains to be meticulously studied.

delta plain began to develop, and at ~4266 cal yr BP, delta plain deposits began to appear, corresponding to high stand system tract (HST). Thus, sea-level change was the major factor controlling the evolution of the sedimentary environment of ZK3. 4.2. Carbon burial records The percentages of OC, TC, IC, and BD in core ZK3 averaged 0.34%, 0.49%, 0.83%, and 1.92 g/cm3, respectively (Fig. 3). The contents of OC, TC, IC, and BD varied irregularly downcore, with larger variability occurring in the overall core. None of the depth profiles for OC content showed the typical exponential decay trend expected under steady-state conditions of OC accumulation and decomposition (Berner, 1980). The percentage of OC was highest in the delta deposit (U1) and limnetic deposit (U4) (average: 0.41%), where primary productivity was high, and the autochthonous OC was relatively high. In contract, the littoral deposit (U3) and fluvial deposit (U5) contained the lowest percentage of OC (average: 0.23%). In the upper core segments (U1), where more marine conditions were identified from top to bottom, sediments accumulated allochthonous marine organic matter, including refractory organic matter transported from elsewhere; high concentrations of suspended particulate matter and repeated cycles of sedimentation and resuspension over diurnal and neap-spring tidal cycles promote high rates of organic matter remineralization within such zones (Abril et al., 1999). OC is critical to improving soil structure and nutrient composition (Nair et al., 2001). The average content of OC in the LHD is less than 4% (Ye et al., 2016, 2018) (range: 0.02–2%), 5–80 times lower than the OC contents in the wetlands of central Florida in the United States (Nair et al., 2001), but twice the OC contents in the Yellow River delta in China (Zhao et al., 2015). The IC content of Core ZK3 sediments varied between 0.01% and 1.83%, with larger variability in the middle section (Fig. 3). The average IC and TC contents peaked in the shallow-sea deposit (U2), where they reached 0.92% and 1.28%, respectively. IC depositions take place largely in shallow-sea settings, mainly because the high redox conditions in the water columns of such marine systems greatly promote decomposition of organic matter (Zhao et al., 2015). The BD was greater than 1.5 in all five sedimentary units of the LHD (Fig. 3) and comparable to the BD in the Yellow River delta (Zhao et al., 2015). The fluvial deposit (U2) before transgression had the highest BD, up to 2 g/cm3, whereas the BD of the delta deposit (U1) was the lowest,

4.3. Carbon sequestration in LHD sediments Apparent sedimentation rates were calculated on the basis of five dating datasets from samples ZK3S1, ZK3S25, ZK3S3, ZK3S27, and OSL1 of core ZK3 (Table 1). In the case of AMS14C dating, we assumed that intact shells were autochthonous and presumably provided more reliable dates than damaged shells, which were more likely to have come from allochthonous sources. Furthermore, wave action and other forms of turbulence could easily redistribute foraminiferal shells. If a younger shell was found in a lower stratum and an older shell in an upper stratum, we used the younger age because relatively young shells are least likely to have been redistributed into old strata (Xue, 2014). The most reliable and reasonable dates were therefore used for the ASRs. The model ASRs were 1.842, 0.184, 0.013, 0.083, 0.084, and 0.084 cm/ yr for depth intervals of 0–3.5, 3.5–11.95, 11.95–12.83, 12.83–16.05, 16.05–23.75, and 23.75–36.7 m, respectively. The ASRs of the eight depositional units were inferred from the above calculation (Table 2). The ASR during U1 was the greatest, approximately 0.25 cm/yr; the ASR during U3 was the lowest, about 0.017 cm/yr, probably because the river course shifted, and ZK3 was far from the estuary at that time. Limnetic deposit (U4) belonged to an isolated, lentic environment with little disturbance and sediment input; the ASR during U4 was therefore 6

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Fig. 4. Scatter plots showing the correlations between the carbon and the clay content. Table 2 Vertical apparent sedimentation rates and apparent mass accumulation rate of TC, OC and IC in different sedimentary environments in LHD. Unit

(U1) 1

Deposition thickness (m) ASR (cm/a) TC-AMAR (g/(m2·yr)) OC-AMAR (g/(m2·yr)) IC-AMAR (g/(m2·yr))

2

3

4

(U1 )

(U1 )

(U1 )

(U1 )

Average

2.25 1.84 200.24 109.12 91.11

2.45 0.94 59.83 34.94 24.89

5.5 0.18 22.58 14.94 7.63

0.8 0.18 29.66 20.30 9.35

11 0.25 52.05 30.77 21.28

TC OC ASR BD TC-AMAR OC-AMAR

1

OC

ASR **

.713 1

-.188 -.060 1

BD *

**

-.444 -.663** -.069 1

TC-AMAR

OC-AMAR

.152 .211** .780** -.246** 1

.078 .269** .780** -.300** .978** 1

*

(U3)

(U4)

(U5)

0.8 0.184 46.46 13.04 33.42

0.6 0.017 26.79 8.44 18.35

11 0.069 18.73 6.01 12.72

13.3 0.084 9.89 3.73 6.16

the OC-AMAR of the delta deposit was lower than that of Florida wetlands (320 g/(m2·yr)) and Louisiana Barataria Basin coastal marshes (300 g/(m2·a)) in the USA as well as some other reported wetlands ((100–200) g/(m2·a)) (Bridgham et al., 2006a; Chmura et al., 2003; Craft, 2007; Turner et al., 2000). However, it was higher than that of the East China Sea shelf (14.7 g/(m2·yr)) (Deng, 2005), Qinghai-Tibet Plateau lakes (19.7 g/(m2·yr)) and Yunnan-Guizhou Plateau lakes (16.6 g/(m2·yr)) (Zhang et al., 2013). Deng (2005) found a similar result: the estuary of the Yangtze River had a higher OC-AMAR than the adjacent continental shelf. The carbon sequestration capacity of the sediments of estuarine delta wetlands is therefore greater than that of continental shelves and plateau lakes.

Table 3 Pearson correlation coefficient matrix of TC, OC, ASR, BD, TC-AMAR and OCAMAR. TC

(U2)

**Correlation is significant at the 0.01 level. * Correlation is significant at the 0.05 level.

4.4. Analysis of factors influencing carbon burial

low, 0.069 cm/yr. Based on the ASRs, the model ages of OC, TC, and IC could be obtained through linear interpolation or extrapolation. We could therefore determine how these parameters changed with time and during the evolution of the different depositional environments. The depth profiles of OC-AMAR, TC-AMAR, and IC-AMAR differed in ZK3. Both the OCAMAR and TC-AMAR gradually decreased from the top to the bottom. The highest rates occurred in U1, 30.77 g/(m2·yr) and 52.05 g/(m2·yr), respectively (Table 2). In contrast, the OC-AMAR of the delta deposit was significantly lower in the LHD (30.77 g/(m2·yr)) than in the Yellow River Delta (1331 g/(m2·yr)) (Zhao et al., 2015), although the OC content in the sediments was lower in Yellow River Delta. Moreover,

As shown in Table 3, both the OC-AMARs and TC-AMARs were significantly correlated with the ASRs (r2 = 0.78, p < 0.01). The indication is that the ASR accounts for ~78% contribution to the rate of carbon sequestration, therefore is the dominant factor controlling changes of TC-AMAR and OC-AMAR, a conclusion also reached by RuizFernández et al. (2018) based on a study of a tropical salt march. However, OC content only explains ~27% of the OC-AMAR (r2 = 0.269, p < 0.01). The OC content is therefore the second important factor that influences OC-AMAR. Baumgart et al. (2010) have shown that the reduction of the residence time of organic matter in oxic surface sediments when the sedimentation rates are high leads to better 7

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Fig. 5. Sediment proxy records in core ZK3. (a) the Northern Hemisphere isolation (Laskar et al., 2004); (b) GISP2-δ18O (Rasmussen et al., 2014); (c) CIA; (d) OCAMAR; (e) age –elevation model.

preservation of the organic matter. Nevertheless, ASRs were weakly correlated with TC and OC in our study (Table 3). This weak correlation may have several causes, including spatial variability of organic matter content and variability of the decomposition rates of organic matter in different sedimentary environments. The Northern Hemisphere insolation (Laskar et al., 2004) and Greenland ice core (GISP2) δ18O record (Rasmussen et al., 2014; Seierstad et al., 2014) were used to relate the response of the carbon burial record to paleoclimatic changes. We therefore focused on the influence of paleoclimatic changes on the OC content and OC-AMR of the sediments and then examined the effect of these changes on carbon burial. The times when the OC content was relatively low value were ~30,000 cal yr BP (U5), 24,500–27,500 cal yr BP (U5), 16,000–17,000 cal yr BP (U4), and 12,000–13,000 cal yr BP (U4). These time intervals correspond to four cold events, namely the Heinrich 3 event (H3), Heinrich 2 event (H2), Heinrich 1 event (H1), and Younger Dryas event (Heinrich, 1988; Hemming, 2004; Rasmussen et al., 2003). These low OC contents were about 30–50% lower than the average OC contents around them. Because ZK3 was far from the North Atlantic icerafting zone, these low OC values were probably caused by changes in the strength of winter monsoon winds over China (Porter and Zhisheng, 1995; Wang et al., 2001). Enhanced westerly winter monsoon winds stimulated by cold North Atlantic sea surface temperatures during Heinrich events could have a dramatic effect on regions downwind, including the LHD, and could reduce OC concentrations by lowering productivity in aquatic systems. Previous studies have shown that the 8.2-kyr BP event was also an important cold event in recorded history (Barber et al., 1999; Pross et al., 2009; Rohling and Pälike, 2005) and

corresponded to the U3 deposition unit in ZK3. However the ASR during U3 was the lowest throughout the core, only 0.017cm/a. Because of the low ASR and absence of high-resolution sampling in this section, the 8.2-kyr event was not well reflected in the test results. In addition, it is apparent from Fig. 5a and c that the OC contents in the sediments of U4 generally follow the gradual increase of Northern Hemisphere summer insolation. The indication is that summer light intensity to some extent controlled the distribution of OC in the Littoral deposit during the interval ~24,000–8,500 cal yr BP. The chemical index of alteration (CIA) is the molar ratio of Al2O3 to Al2O3+CaO*+Na2O + K2O, with CaO* referring to CaO in silicate. The CIA is used to reconstruct chemical weathering (Nesbitt et al., 1982) and Quaternary palaeoclimate changes in East Asia (Chen et al., 2001; Yang et al., 2004a, 2004b). As shown in Fig. 5, CIA values were within the range 43–61, which is in agreement with the H3, H2, H1, and YD events. In addition to the good correspondence with these cold events, the CIA in core ZK3 also demonstrated that there were likely some extra short periods of cold climate during geological history, such as ~500 cal yr BP, ~1000 cal yr BP, ~1500 cal yr BP, and ~4700 cal yr BP, when OC-AMARs were also low (Fig. 5). The two distinct periods that occurred at ~500 cal yr BP and ~1000 cal yr BP in ZK3 probably corresponded to the Little Ice Age (LIA) and are also recognizable in the mud area on the inner shelf of the East China Sea (Liu et al., 2010). The 550-yr BP cooling period was probably the coldest period over the past 2000 years (Xiang et al., 2006). According to historical records (Hong et al., 2000; Jin et al., 2003; Liu et al., 1999; Shi et al., 1999; Wang et al., 1998; Yao et al., 1990), the LIA had extensive imprints around the world: in Europe, North America, Greenland, Antarctica, the Dunde 8

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ice core, a stalagmite from Shihua Cave, and in peat from Jinchuan and Hongyuan. Furthermore, the grain size proxies from core B2 in the mud area southwest of Cheju Island also show obvious climate fluctuations during the LIA and record three distinct cooling events. The results obtained from Core ZK3 are therefore consistent with regional changes during the LIA. Meanwhile, the ~1500 cal yr BP and ~4700 cal yr BP cold events from the CIA proxy have also been found in the Zhujiang (Pearl River) estuary as several abrupt decreases of OC concentrations (Yang et al., 2011) that are related with ice-rafted debris events in the North Atlantic (Bond et al., 2001). We conclude from the foregoing analysis that cold climate events during geological history in the LHD affected OC preservation in the sediments and controlled the carbon burial efficiency of the sediments.

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5. Conclusion (1) The contents of OC in the whole core ranged from 0.02% to 2%. The OC contents of the delta deposit and limnetic deposit were the highest in terms of high productivity, and those of the littoral deposit and fluvial deposits were the lowest. Both the average IC and TC contents peaked in the shallow-sea deposit, where they reached 0.92% and 1.28%, respectively. Sediment grain size was one factor that controlled C abundance. (2) The ASRs of the units varied between 0.069 cm/yr and 0.25 cm/yr in ZK3. The OC-AMAR gradually decreased from the top to the bottom of the core. The highest rate occurred at the delta deposit, 30.77 g/(m2·yr). The carbon sequestration rate was higher in the sediments of the estuarine deltaic wetlands than in continental shelf sediments. (3) The ASR was the dominant modulator of the OC-AMARs and TCAMARs, which were, however, poorly correlated with BD. The OCAMARs were significantly (p < 0.01) correlated with OC contents. The CIA proxy in ZK3 recorded several cold events during the last ~40,000 years. The good correspondence of these events with low OC-AMAR indicated that cold climate events during geological history in the LHD adversely affected carbon burial in the sediments. Acknowledgements This study was jointly funded by the Natural Science Foundations of China (Grant No. 41406082, 41240022, 41706057), National Key R&D Program of China (2016YFE0109600), and Governmental Public Research Funds of China (No.201111023, DD20189503 and GZH201200503). We would like to thank Prof. Chunting Xue for his help in sedimentation analysis. References Abrahim, G.M., Nichol, S.L., Parker, R.J., Gregory, M.R., 2008. Facies depositional setting, mineral maturity and sequence stratigraphy of a Holocene drowned valley, Tamaki Estuary, New Zealand. Estuar. Coast Shelf Sci. 79, 133–142. Abril, G., Etcheber, H., Le Hir, P., Bassoullet, P., Boutier, B., Frankignoulle, M., 1999. Oxic/anoxic oscillations and organic carbon mineralization in an estuarine maximum turbidity zone (The Gironde, France). Limnol. Oceanogr. 44, 1304–1315. Anbuselvan, N., Senthil Nathan, D., Sridharan, M., 2018. Heavy metal assessment in surface sediments off Coromandel Coast of India: implication on marine pollution. Mar. Pollut. Bull. 131, 712–726. Barber, D.C., Dyke, A., Hillaire-Marcel, C., Jennings, A.E., Andrews, J.T., Kerwin, M.W., Bilodeau, G., McNeely, R., Southon, J., Morehead, M.D., 1999. Forcing of the cold event of 8,200 years ago by catastrophic drainage of Laurentide lakes. Nature 400, 344. Baumgart, A., Jennerjahn, T., Mohtadi, M., Hebbeln, D., 2010. Distribution and burial of organic carbon in sediments from the Indian Ocean upwelling region off Java and Sumatra, Indonesia. Deep Sea Res. Oceanogr. Res. Pap. 57, 458–467. Berner, R.A., 1980. Early Diagenesis: A Theoretical Approach. Princeton University Press, Princeton, USA, pp. 245. Bond, G., Kromer, B., Beer, J., Muscheler, R., Evans, M.N., Showers, W., Hoffmann, S., Lottibond, R., Hajdas, I., Bonani, G., 2001. Persistent solar influence on North Atlantic climate during the Holocene. Science 294, 2130–2136. Bridgham, S., Megonigal, J.P., Keller, J., Bliss, N., Trettin, C., 2006a. The carbon balance

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