Sedimentary Evolution of the Holocene Subaqueous Clinoform off the Southern Shandong Peninsula in the Western South Yellow Sea

QIU Jiandong, LIU Jian,, SAITO Yoshiki, YANG Zigeng, YUE Baojing, WANG Hong, and KONG Xianghuai

1) College of Marine Geo-Science, Ocean University of China, Qingdao 266100, P. R. China

2) The Key Laboratory of Marine Hydrocarbon Resources and Environment Geology, Ministry of Land and Resources, Qingdao 266071, P. R. China

3) Qingdao Institute of Marine Geology, Ministry of Land and Resources, Qingdao 266071, P. R. China

4) Geological Survey of Japan, AIST, Central 7, Higashi 1-1-1, Tsukuba, Ibaraki 305-8567, Japan

Sedimentary Evolution of the Holocene Subaqueous Clinoform off the Southern Shandong Peninsula in the Western South Yellow Sea

QIU Jiandong1),2),3), LIU Jian2),3),*, SAITO Yoshiki4), YANG Zigeng3), YUE Baojing3), WANG Hong3), and KONG Xianghuai3)

1) College of Marine Geo-Science, Ocean University of China, Qingdao 266100, P. R. China

2) The Key Laboratory of Marine Hydrocarbon Resources and Environment Geology, Ministry of Land and Resources, Qingdao 266071, P. R. China

3) Qingdao Institute of Marine Geology, Ministry of Land and Resources, Qingdao 266071, P. R. China

4) Geological Survey of Japan, AIST, Central 7, Higashi 1-1-1, Tsukuba, Ibaraki 305-8567, Japan

Based on the stratigraphic sequence formed since the last glaciation and revealed by 3000 km long high-resolution shallow seismic profiles and the core QDZ03 acquired recently off the southern Shandong Peninsula, we addressed the sedimentary characteristics of a Holocene subaqueous clinoform in this paper. Integrated analyses were made on the core QDZ03, including sedimentary facies, sediment grain sizes, clay minerals, geochemistry, micro paleontology, and AMS14C dating. The result indicates that there exists a Holocene subaqueous clinoform, whose bottom boundary generally lies at 15–40 m below the present sea level with its depth contours roughly parallel to the coast and getting deeper seawards. The maximum thickness of the clinoform is up to 22.5 m on the coast side, and the thickness contours generally spread in a banded way along the coastline and becomes thinner towards the sea. At the mouths of some bays along the coast, the clinoform stretches in the shape of a fan and its thickness is evidently larger than that of the surrounding sediments. This clinoform came into being in the early Holocene (about 11.2 cal kyr BP) and can be divided into the lower and upper depositional units (DU 2 and DU 1, respectively). The unit DU 2, being usually less than 3 m in thickness and formed under a low sedimentation rate, is located between the bottom boundary and the Holocene maximum flooding surface (MFS), and represents the sediment of a post-glacial transgressive systems tract; whereas the unit DU 1, the main body of the clinoform, sits on the MFS, belonging to the sediment of a highstand systems tract from middle Holocene (about 7–6 cal kyr BP) to the present. The provenance of the clinoform differs from that of the typical sediments of the Yellow River and can be considered as the results of the joint contribution from both the Yellow River and the proximal coastal sediments of the Shandong Peninsula, as evidenced by the sediment geochemistry of the core. As is controlled mainly by coactions of multiple factors such as the Holocene sea-level changes, sediment supplies and coastal dynamic conditions, the development of the clinoform is genetically related with the synchronous clinoform or subaqueous deltas around the northeastern Shandong Peninsula and in the northern South Yellow Sea in the spatial distribution and sediment provenance, as previously reported, with all of them being formed from the initial stage of the Holocene up to the present.

subaqueous clinoform; Holocene; Yellow Sea; Shandong Peninsula; Yellow River; provenance; sea-level change; sedimentary

1 Introduction

Under the strong wave and tidal actions in coastal zones, sediments from rivers form river deltas at their mouths and are transported offshore resulting in building up clinoforms or subaqueous deltas in coastal and shelf areas. These subaqueous deltas are mostly composed of muddy sediments. In China, such clinoforms more than 100 km apart from the river mouth have been found in the areas off the Shandong Peninsula in the Yellow Sea due to the sediment supply from the Yellow River (Cheng et al., 2001; Liu et al., 2004; Liu et al., 2007b), in the coastal areas of Fujian and Zhejiang Provinces in the East China Sea with the sediments from the Yangtze River (Liu et al., 2007a; Shi et al., 2010) and in the areas off the Guangdong coast in the South China Sea due to the sediment input from the Pearl River (Ying, 1999). Similar alongshore sediment transport and clinoforms are reported also existing due to other large rivers: the Mekong River (Ta et al., 2002a, b), Amazon River (Nittrouer et al., 1996), Po River (Cattaneo et al., 2003), and Fly River (into Gulf of Papua New Guinea) (Walsh et al., 2004; Slingerlandet al., 2008).

Clinoforms around the Shandong Peninsula in the Yellow Sea are typical offshore clinoforms or offshore subaqueous deltas. Those in the South Yellow Sea is named Shandong subaqueous delta (Alexander et al., 1991). Major sediment source of these clinoforms is the Yellow River, which is the second largest river in the world in terms of sediment discharge. Its mean annual sediment discharge into sea is about 1.1×109t (Milliman and Meade, 1983). Sediments transported from the Yellow River are mostly fine-grained, and ～95% of sediments consists of silt and clay (Qin et al., 1990). The river forms a huge delta at its mouth in the Bohai Sea during the Holocene (Saito et al., 2000), and dispersed fine-grained sediments are transported offshore passing the Bohai Strait into the Yellow Sea along the Shandong Peninsula. These fine-grained sediments are one of the provenances of the muddy sediments around the Shandong Peninsula in the Yellow Sea, which build up a subaqueous clinoform (Qin and Li, 1986; Milliman et al., 1986; Alexander et al., 1991; Martin et al., 1993; Li et al., 2004; Liu et al., 2004; Yang and Liu, 2007; Liu et al., 2007a,b; Wang and Li, 2009). The estimated quantity of the Yellow River sediment transported outwards into the Yellow Sea ranges from 1% (Martin et al., 1993) to 15% of the total (Alexander et al., 1991). Recent studies show the detailed isopach of the Holocene sediments and tongue-like shape of the subaqueous delta off the southern Shandong Peninsula (Yang and Liu, 2007) (Fig.1) and the stratigraphy and environmental changes during the Holocene off the northern Shandong Peninsula (Liu et al., 2007b). However, still we do not have enough data for the main part of the Shandong subaqueous delta. In this paper, we discuss sedimentary evolution of the clinoform located south of the Shandong Peninsula in the Yellow Sea, which is a western marginal part of the Shandong subaqueous delta. Particularly, stratigraphy, ages, and sediment source of the clinoform are dealt with based on the analyses of high-resolution shallow seismic profiles covering approximately 3000 km and a 40.20-m long core taken in the study area. These results will be of great significance to further understanding of the transportation of the Yellow River sediments into the Yellow Sea and the spatial distribution of the subaqueous delta.

Fig.1 Bathymetry and regional circulation pattern in the Yellow Sea and adjacent areas during winter (modified after Su and Yuan, 2004). The shaded area around the Shandong Peninsula denotes the distribution of a subaqueous clinoform (Yang and Liu, 2007). Water depth and the thickness of the clinoform are in meters. The dotted square indicates the study area. SYS, South Yellow Sea; NYS, North Yellow Sea; BS, Bohai Strait; KC, Kuroshio Current; YSWC, Yellow Sea Warm Current; TC, Tsushima Current; TWC, Taiwan Warm Current; YSCC, Yellow Sea Coastal Current; SSCC, South Shandong Coastal Current; SKCC, South Korean Coastal Current; NJCC, North Jiangsu Coastal Current; ECSCC, East China Sea Coastal Current; JTSR, Jiangsu tidal sand ridges.

2 Regional Setting

2.1 Geology

Bedrock of the Shandong Peninsula consists of a basement of Archaean–Proterozoic crystalline metamorphic rocks overlain by Neoproterozoic–Sinian shallow-marine sandstone and limestone, Jurassic–Neogene terrestrial red clastic sediments intercalated with intermediate-basic and intermediate-acidic volcanic rocks, and Neogene basalt. Overlying Quaternary sediments are widely distributed, with unconsolidated fluvial–lacustrine, proluvial, and eluvial sediments as the dominant facies (Li et al., 1994).

The South Yellow Sea (SYS) is a large Mesozoic and Cenozoic composite basin consisting of uplifts and depressions. The tectonic evolution of the SYS can be divided into four principal stages: the Paleozoic-Triassic marine basin, the Mesozoic land basin, the Paleocene-Middle Miocene fault basin, and the regional depression since the late Miocene (Yao et al., 2005).

2.2 Topography and Rivers in the Shandong Peninsula

The Shandong Peninsula has an indented coastline with many bays (Fig.2). The Rushan and Juhe Rivers contribute an annual sediment discharge of (30–50) ×104t to Rushan Bay. The annual sediment discharge of the Wulong River (124 km long) averages about 84×104t to Dingzi Bay (data from 1958–1980). Seasonal short streams contribute small annual sediment discharges of (2.31–3.90) ×104t to Aoshan and Xiaodao bays. The Dagu River is the largest fluvial system that discharges into Jiaozhou Bay. Its main channel is 180 km long, and its historical annual sediment discharge into the bay is 95.9×104t. Since the late 1970s, however, its middle and lower reaches have run dry except during flood seasons (Editorial Board of Annals of Bays in China, 1993). The Yellow River presently enters the Bohai Sea, and the sediments from the river can be transported into the North and South Yellow Sea by alongshore currents. The Yellow River has changed its course frequently during the Holocene, and a major channel shift in the last 1000 years took place from 1128 to 1855 AD, when the river shifted south of the Shandong Peninsula and emptied into the SYS (Zhang, 1984).

The Yellow Sea is a broad, relatively shallow, and low gradient epicontinental sea above a flat, tectonically stable shelf (Fig.1). Water depth in most parts of the South Yellow Sea is less than 100 m. The SYS has a maximum width of about 700 km and deepens toward the Yellow Sea Trough, defined by an 80-m isobath that extends NWSE (Qin et al., 1989b).

The floor of the western SYS is made up of four geomorphologic units from north to south: 1) a Holocene mud clinoform around the tip of the Shandong Peninsula called the Shandong mud wedge (Liu et al., 2004, 2007b); 2) an erosional area south of the Shandong Peninsula that consists of a mixture of gravel, sand, and mud (Milliman et al., 1987); 3) the Old Yellow River Delta, an abandoned subaqueous delta formed between 1128 and 1855; and 4) the Jiangsu tidal sand ridges, large-scale radiating tidal sand ridges produced by radial tidal currents.

Fig.2 Map of the study area showing distribution of high-resolution shallow seismic profiles and locations of core QDZ03 and the profiles in Figs.3 and 4.

2.3 Ocean Currents and Shelf Sediments

Several current systems are active in the Yellow Sea and adjacent regions (Fig.1). After entering the Yellow Sea from the southwest side of Cheju Island, the Yellow Sea Warm Current gradually turns northward and runs along the west side of the Yellow Sea deep trough. It turns west upon entering the North Yellow Sea and may extend into the Bohai Sea in winter, where its salinity drops below 32. The East Yellow Sea coastal current, also called the South Korean Coastal Current, flows southward in winter and northward in summer roughly along the 20–40 m depth contours on the west side of the Korean Peninsula. The coastal currents (the Yellow Sea, South Shan-dong, and North Jiangsu Coastal Currents) along the western coast of the Yellow Sea flow persistently southward during both summer and winter (Beardsley et al., 1983). The West Yellow Sea Coastal Current, joined by the Bohai current system, runs eastward along the north coast of the Shandong Peninsula, makes a turn around Chengshan Cape, and then flows southward out of Haizhou Bay at depths of 40–50 m. It changes direction to the southeast near 33˚–32˚N, with its front approaching 30˚N. Although the route of the coastal currents in the Western Yellow Sea remains unchanged throughout the year, their flux and velocities have evident regional differences (Su and Yuan, 2004).

Although most of the sediment of the Yellow River is deposited in the Bohai Sea, a small portion of the sediment is carried by the coastal currents to the nearshore areas and accumulates along the south and north coasts of the Shandong Peninsula (Liu et al., 2007b; Yang and Liu, 2007). The surface sediment of the study area is dominated by clayey silt (Li et al., 2005). Previous studies of its detrital mineralogy, clay mineral composition, and geochemistry (Chen et al., 1986; Li et al., 2000; Li et al., 2010) and of the horizontal distribution of suspended matter in the SYS (Qin et al., 1989a) have shown that the surface sediment came largely from the Yellow River.

3 Data and Methods

From May to July 2009, the Qingdao Institute of Marine Geology measured 3000 km of high-resolution shallow seismic profiles in the study area (Fig.2) to elucidate the shallow stratigraphy of the nearshore area off the southern Shandong Peninsula. The survey comprised 28 east-west transects with a spacing of 2.5 km and 23 north–south transects with a spacing of 5 km (Fig.2). The seismic survey used an SBP/AAE high-resolution profiler (Applied Acoustic Engineering Company, UK) with an excitation energy of 300 J, excitation interval of 480 ms, and band-pass filtering of 320–2000 Hz.

Depositional sequences and bounding surfaces (onlap, downlap, truncation, and toplap) were identified from the character of reflectors in the seismic records. We selected an acoustic velocity of 1550 m s-1for subbottom strata based on previous drilling and seismic experience.

Core QDZ03 with a length of 40.20 m was obtained by continuous drilling in May 2011 by drilling vessel Kan 407, operated by the Shanghai Offshore Petroleum Geophysical Corporation. The core site (36˚16΄03.21΄΄N, 120˚56΄58.77΄΄E) is located in muddy sediment at a water depth of 15.50 m. Recovery of the mud-dominated sediments of the core was greater than 87% on average, and recovery of the sandy intervals was generally 60% to 90%.

The core was halved, photographed, described, and sampled in the laboratory. Sampling was done at intervals of 10–12 cm for clay mineralogy and grain size analysis and 10–15 cm for micropaleontologic identification, each sample being no thicker than 4 cm.

Benthic foraminifers and unbroken shells were selected and sent to Beta Analytic (Miami, USA) for accelerator mass spectrometry (AMS)14C dating. Eight ages were obtained (Table 1).14C ages were corrected for the regional marine reservoir effect (ΔR=-100±36 yr) determined for the western and eastern Yellow Sea afterSouthon et al. (2002), and calibrated using Calib Rev. 5.0.1 (Stuiver et al., 2005). Calibrated ages are given with one standard deviation (1σ) uncertainty (Table 1). Calibrated ages in this paper are reported as years before AD 1950 (cal yr BP), and uncalibrated ages are given as14C years BP (14C yr BP).

Table 1 AMS14C dating results from core QDZ03

We took nine bottom sediment samples from small rivers that discharge into the coastal bays (seven from the Wulong River and two from the Rushan River) within 30 km upstream from the bay areas to examine their geochemistry.

Analyses of clay minerals, grain size, micropaleontology (foraminifers and ostracods), and rare earth elements (REEs) for the sediments of core QDZ03 and analyses of REEs for the bottom sediment of the small rivers were conducted in the Testing Center of Qingdao Institute of Marine Geology. Before X-ray diffraction analysis of clay minerals, each sample was washed twice with distilled water to remove salt. Organic matter was removed by hydrogen peroxide. Sodium hexametaphosphate at 0.1 mol L-1concentration was added to the samples, which were then kept overnight before extraction of the supernatant liquid, which contained sediment particles smaller than 2 μm. This procedure was repeated at least three times. The sediment in the supernatant liquid was concentrated by centrifugation and then smeared evenly on glass slides, which were dried in air and then kept in an ethylene glycol vapor bath for 24 h. Clay mineralogy was identified using a D/Max-2500 X-ray diffractometer (Cu target, tube voltage 40 kV, tube current 150 mA).

About 10 g of the sediment was taken for grain-size analysis. Samples were soaked in hydrogen peroxide and dilute hydrochloric acid to remove both organic matter and carbonate, then washed to remove salt. After dispersal in sodium hexametaphosphate solution by ultrasonication, samples were analyzed for grain size using a Mastersizer-2000 laser particle analyzer (Malvern Company, UK) with a measuring range of 0.02–2000 μm, particle resolution of 0.01Φ, and relative error of <2%. Micropaleontologic identifications were made with standard methods (Wang et al., 1985). REEs of the bulk sediment were determined by inductively coupled plasma mass spectrometry (ICPMS) with the method reported by Jiang et al. (2011).

Thirteen samples for optically stimulated luminescence (OSL) dating were taken below the core depth of 13.0 m. OSL ages were determined in the Testing Center of the Qingdao Institute of Marine Geology with a Daybreak 2200 TL/OSL reader. All samples showed high OSL sensitivity and a stable OSL signal. Because replicate De values were unscattered, unimodal, and normally distributed, final De values could be derived with the Central Age Model (Galbraith et al., 1999). Concentrations of K, U, Th, and Rb were measured by ICP-MS and were converted into dose rates based on data from Adamiec and Aitken (1998) and Marsh et al. (2002). Cosmic dose rate was estimated according to Prescott and Hutton (1994). OSL ages were reported relative to AD 2010. Results of OSL dating are listed in Table 2.

Table 2 OSL dates from core QDZ03

4 Results

4.1 Seismic Stratigraphy

The shallow seismic profile data allowed us to divide the sediments between the seafloor (seismic reflector T0) and the top of the bedrock in the study area into five seismic stratigraphic units (named SU 1–SU 5 from top to bottom) separated by four principal bounding surfaces (T1–T4) (Figs.3–5). The unit above T1 is related to the clinoform that is the subject of our study.

Seismic reflector T3 is an erosional surface of high relief that includes V- or U-shaped incisions in the underlying strata (Figs.3–5). Compared with the other seismic reflectors, T3 has larger undulations and is more discontinuous laterally. The incised valleys change greatly in depth and width within the study area. Some large trunk valleys are deeper than 15 m and as wide as 2 km. Unit SU 3 at the bottom of the incised valleys is generally <10 m thick and is characterized by chaotic or wavy reflections of high amplitude, which may be cut by the overlying surface T2.

T2 is a wavy erosional surface of high amplitude, which slightly incises the underlying SU 3. Unit SU 2, between T2 and T1, shows a horizontal-to-wavy reflector of high amplitude, locally with incised tidal channels and chaotic reflectors. The thickness of the unit varies from <1 to 10 m. T2 extends to 40–50 m below present sea level (bpsl)in some of the incised valleys.

Fig.3 Typical N-S shallow seismic profile (location shown in Fig.2). The red broken line above T1 represents the maximum flooding surface of the post-glacial period. SU 1–SU 5 denote seismic units; T1–T4 correspond to the bounding surfaces between SU 1–SU 5, respectively, and T0 to seafloor; TWTT, two-way travel time; MFS, maximum flooding surface.

T4 is a relatively smooth erosional surface that dissects the underlying layers and locally fluctuates slightly. Unit SU 4 between T3 and T4 changes drastically in thickness because of the downcutting of T3. With a maximum thickness of 15.5 m, SU 4 is characterized by parallel or subparallel high-amplitude reflections that dip seaward locally. It becomes thinner toward land and pinches out in some places. Unit SU 5 between T4 and the basement surface shows a wavy and chaotic reflection pattern, which is difficult to trace because of the interference of multiple reflecting interfaces. The top of seismic basement shows strong reflections, and its depth ranges from 0 to >100 m in the study area. In general, it deepens rapidly from the coast to the shallow sea, but becomes shallow around islands and is exposed at the bottom of tidal channels near some islands and the outlet of Jiaozhou Bay.

Fig.4 Typical W-E shallow seismic profile (location shown in Fig.2). The red broken line above T1 represents the maximum flooding surface of the post-glacial period. SU 1–SU 5 denote seismic units; T1–T4 correspond to the bounding surfaces between SU 1–SU 5, respectively, and T0 to seafloor; TWTT, two-way travel time; MFS, maximum flooding surface.

Fig.5 Typical W-E shallow seismic profile (location shown in Fig.2). The red broken line above T1 represents the maximum flooding surface of the post-glacial period. SU 1–SU 5 denote seismic units; T1–T4 correspond to the bounding surfaces between SU 1–SU 5, respectively, and T0 to seafloor; TWTT, two-way travel time; MFS, maximum flooding surface.

4.2 Distribution of the Clinoform

Reflector T1, the base of the clinoform, is a gently undulating erosional surface that dissects the underlying unit SU 2. It is characterized by an intermediate-to-strong amplitude and can be traced continuously in the study area. It is usually at depths of 15–40 m bpsl (Fig.6). Unit SU 1 lies between T1 and T0 and is divided here into a lower and an upper unit. Its lower unit (mostly <3 m thick) exhibits basal onlap and is overlaid by accretionary subhorizontal reflectors. The upper unit has gently inclined reflectors that prograde to the south and east. We construed the boundary between the lower and upper units of SU 1 to be the maximum flooding surface (MFS) of the post-glacial period (Figs.3–5).

Generally, SU 1 is a subaqueous clinoform that progrades southward and eastward; its thickness distribution is shown in Fig.7. In the study area, it extends parallel to the coastline and is thickest to the north of Laoshan Head, thinning seaward. In the mouths of Xiaodao, Aoshan, Dingzi, and Rushan bays, it forms fans with greater thicknesses than in the adjoining areas. Its greatest thickness, 22.5 m, is at the mouths of Aoshan and Xiaodao bays. The 3-m isopach appears to mark a boundary; landward the unit’s thickness varies greatly, and it forms a steep slope; seaward its thickness varies from 1 to 3 m, and it forms relatively smooth slopes. The fact that the 3-m isopach runs roughly along the present 25-m depth contour shows that the submarine clinoform of the southern Shandong Peninsula is mainly distributed above a water depth of 25 m along the coast , which is is consistent with previous surveys (Qin et al., 1989b). A subaqueous clinoform, no doubt, exists off the southern Shandong Peninsula.

Fig.6 Depth distribution of seismic reflector T1 in meters below present sea level (bpsl) in the study area.

Fig.7 Thickness distribution of sedimentary sequences above seismic reflector T1 in the study area.

4.3 Lithology and Ages of Core QDZ03

Core QDZ03 is 40.20 m long and penetrates the clinoform and a small incised valley (Fig.3). This paper presents details of the top 12 m of the core as well as OSL dates from the entire core to correlate with the seismic units and global marine isotope stages (Fig.8). We used comparisons of the core with the shallow seismic profile across the core site to divide the stratigraphic sequence from 0 to 12 m into three depositional units named DU1, DU 2, and DU 3, in descending order.

DU 3 (8.63–12.00 m) corresponds to the upper part of seismic unit SU 2. It consists mainly of brown, brownyellow, and dark gray clayey silt to fine sand with clayey bands. Brown rust patches and black carbonaceous spots are common. The foraminiferal assemblage is dominated by the euryhaline, nearshore, shallow-water species Ammonia beccarii (Linné) vars. and subordinate nearshoreinner shelf common species Elphidium advenum (Cushman) and E. hispidum Cushman. Ostracods are mainly represented by euryhaline continental species such as Candoniella albicans (Brady) and Candoniella spp. and by subordinate coastal shallow-water species Neomonoceratina chenae Zhao et Whatley, Keijella bisanensis (Ocubo), Albileberis sinensis Hou, and Pistocythereis bradyformis (Ishizaki). Some samples contain no foraminifers and ostracods, indicative of a continental environment. Based on the lithology of the core and the reflection pattern in seismic profiles, we construed DU 3 as estuarine-floodplain sediment. The widely developed brown rust spots suggest oxidation from repeated subaerial exposure.

DU 2 (7.70–8.63 m) corresponds to the lower unit of SU 1 showing accretionary or onlapping subhorizontal reflectors. The sharp erosional surface at its base, which corresponds to seismic surface T1, is marked by a 3-cmthick shelly layer with mud pebbles. DU 2 is mainly composed of green and yellow-gray clayey silt intercalated with lenticular beds and thin bands of silt or fine sand. These coarser-grained intercalated bands decreaseupward gradually, showing a normal grading. The mean grain size is 6.0 Φ. Sporadic shell fragments and moderate bioturbation are present. The foraminiferal assemblage is dominated by the nearshore shallow-water species Elphidium magellanicum Heron-Allen et Earland and A. beccarii vars. with subordinate nearshore to inner-shelf common species Protelphidium tuberculatum (d’Orbigny) and Cribrononion subincertum (Asano). Ostracods are dominated by euryhaline nearshore shallow-water species Sinocytheridea impressa (Brady) and N. chenae and to a lesser extent by K. bisanensis and P. bradyformis. Based on its lithologic character, we construed DU 2 as having been deposited in a tide-influenced subtidal nearshore environment. Three AMS14C dates in a good age order show a consistent time of the early Holocene from 10638 to 9336 cal yr BP. Extrapolation from the sedimentation rates calculated with these three dates gives the age range of DU 2 as roughly 11.0–9.0 cal kyr BP.

DU 1 (7.70–0 m) corresponds to the upper unit of SU 1 showing gently inclined reflectors. It is mainly composed of green-gray and yellow-green-gray clayey silt that contain sporadic shell fragments. The lower part (4.80– 7.70 m) has very uniform grain sizes, with mean values from 6.9 to 7.2 Φ (7.0 Φ on average), whereas the sediment above 4.80 m has a wider range in mean grain sizes from 5.8 to 7.2 Φ and a slightly coarser average as a whole (6.6 Φ on average). The basal bounding surface is marked by an abrupt increase in grain-size below it in DU 2. We construed this surface to correspond to seismic surface MFS. The foraminiferal assemblage mainly consists of A. beccarii vars. and subordinate P. tuberculatum, E. magellanicum, and E. advenum; ostracods are dominated by S. impressa and subordinate N. chenae and K. bisanensis. Based on its lithology, we construed DU 1 as having been deposited in nearshore shallow water that was deeper than was the case for DU 2. Five AMS14C dates of DU 1 ranged from 4796 cal yr BP at the bottom to 559 cal yr BP in the upper part, an indication that DU 1 was formed during the middle and late Holocene. There is a hiatus of about 4000 years between DU 1 and the underlying DU 2.

Fig.8 Lithologic column of core QDZ03 with AMS14C dates (a) and distribution of mean grain size (Mz) in the 0–9 m interval (b). DU 1–DU 3: depositional units; T0–T1: seismic bounding surfaces; MFS: Holocene maximum flooding surface.

The core section below T4 (28.9–40.20 m) is characterized by interbedded sand and mud with common brown rusty spots. Its OSL dates ranged from 126 to more than 140 ka BP (Fig.8 and Table 2), and we construed it to be fluvial deposits during Marine Isotope Stage (MIS) 6. The core section between T4 and T3 (28.9–17.4 m) is dominated by clayey silt that is intercalated with thin (<1 cm) layers or lenses of silt to fine sand. Local millimeterto centimeter-scale bands of alternating clayey silt and silt to fine sand that exhibit wavy and lenticular bedding are construed to be shallow-sea deposits influenced by tidesand are associated with MIS 5 on the basis of OSL dates between about 75 and 80 ka (Fig.8 and Table 2) in the middle part of this section. The core section between T3 and T2 (17.4–13.6 m) represents incised-channel infilling above T3 and is largely composed of laminated clayey silts to fine sand with brown rusty stains that are suggestive of subaerial exposure. We construed the incised channel (T3) and the infilling sediments as having formed during MIS 4, in agreement with the OSL dates in the upper part of this section. The core section between T2 and T1 (13.6–8.63 m) consists mainly of interbeds of clayey silt and fine sand with common reddish mottles and brown rusty spots and layers that contain foraminifers in the upper part (see above for detail), and is construed as marine and fluvial sediments deposited during MIS 3 to 2.

4.4 Compositions of Clay Minerals and REEs in the Clinoform

Clay minerals in the 0–8.63 m interval of core QDZ03 consist of illite (46.7%–63.0%), smectite (9.3%–20.2%), chlorite (11.6%–21.4%), and kaolinite (10.4%–17.0%), and the ratio of illite to smectite is generally less than 6 (Fig.9). High percentages of smectite, second only to illite, in the clay mineral assemblages are typical of the sediments from the Yellow River (Yang, 1988).

In plots of chondritenormalized REEs, the pattern in sediments of 0–8.63 m interval has a negative slope, with light REEs having higher ratios than heavy REEs (Fig. 10a). Cerium shows a very weak negative anomaly, the δCe being 0.95–1.00 (averaging 0.98), which is nearly equal to the Ce anomaly (0.97) in the overbank and bottom sediments of the Yellow River (Yang et al., 2003c). Europium is moderately depleted, the δEu being 0.67–0.73 (averaging 0.70), less strong than the Eu anomaly (0.60) in the overbank and bottom sediments of the Yellow River (Yang et al., 2003c). The concentration of REEs (∑REE) ranges between 155.8 and 195.5 μg g-1(Fig.10b), averaging 182.4 μg g-1, which is distinctly higher than the∑REE (148.1 μg g-1) of Yellow River sediments (Yang et al., 2003c).

Fig.9 Depth profile of clay mineral compositions in core QDZ03. DU 1–DU 3: depositional units.

Fig.10 Distribution pattern of chondrite-normalized REE (a) and ∑REE-LREE/HREE diagram (b) of representative samples from 0–9 m interval of core QDZ03, sediment of the Yellow River, granite of Laoshan Mountain, and lamprophyre of Tianheng Island (data for Yellow River sediment, Laoshan granite, and Tianheng lamprophyre are from Yang and Li, 1999; Yang et al., 2003b, 2003c; Han et al., 1991, 2010).

5 Discussion

5.1 Ages of Sedimentary Sequence and Sea-Level Changes

As described in section 4.3, OSR dates show that the core sections below T1, which correspond to SU 2 through the upper part of SU 5, represent sediment deposited during MIS 6 to MIS 2. The core section between T1 and T2 (SU 2) yielded only one OSL date in the lower part, and we infer that this well-stratified unit was formed mostly in a marine environment during MIS 3, inferred by the presence of benthic foraminifers. The uppermost part, however, was deposited during MIS 2 in a fluvial environment, as indicated by the chaotic (opaque) reflections. The brown rusty spots in the core section between T1 and T2 are construed to have resulted from subaerial exposure during the sea-level lowstand of MIS 2.

The core shows that SU 1 was deposited mainly during the Holocene, and its base (bounding surface T1) is construed as a ravinement surface formed during the postglacial sea-level rise. The lower part of SU 1 consists of mainly onlapping and accretionary subhorizontal reflectors, and its thickness is consistently near 3 m. However, the upper part of SU 1 is characterized by inclined reflectors gently downlapping onto the boundary between the upper and lower units. This boundary is construed as a maximum flooding surface (MFS) formed in the mid-Holocene, which is at about 7–6 cal kyr BP in the clinoform northeast of the Shandong Peninsula (Liu et al., 2007b).

The post-glacial sea-level changes in the western Pacific Ocean (Liu et al., 2004; Hanebuth et al., 2011) show a stepwise or episodic pattern (Fig.11), generated by a series of meltwater pulse events associated with climatic fluctuations (Fairbanks, 1989; Bard et al., 1990; Blanchon and Shaw, 1995; Weaver et al., 2003; Tornqvist et al., 2004). The depth of reflector T1 is 40–15 m bpsl in the study area, and its formation time inferred from sea-level change curves is about 11.2–9.0 cal kyr BP. The evidence suggests that the clinoform in the study area began developing in the early Holocene, which coincides with the result of Liu et al. (2007b) that a subaqueous clinoform off the northeast Shandong Peninsula began forming about 11.6 cal kyr BP. The two clinoforms around the Shandong Peninsula are connected in the region. As a ravinement surface moving landward in response to sea-level rise, however, T1 is diachronous, and its formation time should become later from the open sea to the coast. In the western SYS south of the study area, the T1-correlative surface found in cores after the Last Glacial Maximum has been dated at 11–13 cal kyr BP (Liu et al., 2010). Unit DU 2 in core QDZ03 was formed in a subtidal nearshore environment with erosional surface T1 at its base. The age range of DU 2 (about 11.0–9.0 cal kyr BP) is consistent with the formation time of the underlying T1. Most of the transgressive sands have been transported landward or seaward by waves and tides. The fact that some shells in the transgressive sands (DU 2) in core QDZ03 show slightly old ages relative to the sea-level curves (Fig.11) may indicate that they were transported landward.

Our inference of a Holocene MFS as the bounding surface between the lower and upper units of SU 1 is consistent with the lithology of the core sections (DU 1 and DU 2) corresponding to the seismic unit SU 1. DU 2 corresponds to the lower unit of SU 1 and was formed during the early Holocene. The generally fining-upward deposits of DU 2, which match the early sedimentary record of the Holocene clinoform off the northeastern Shandong Peninsula (Liu et al., 2007b), are in accord with the coeval environmental changes accompanying the sea-level rise, which included increasing water depth and a change in depositional setting from proximal to distal. DU 1 represents the sedimentary record during the high sea level since the middle Holocene. The lower part of DU 1 (4.80– 7.70 m in core QDZ03) has very uniform grain sizes, consistent with a stable sea level and coastal environment. In the upper part of DU 1 (0–4.80 m), grain sizes are slightly coarser and less uniform. This change may be related to intensified human activities during the last ～3000 years.

The fact that seismic and depositional units identified in the study area correlate well with those seen in the sedimentary sequence in the coastal area north and east of the Shandong Peninsula (Liu et al., 2007b) highlights the key role of sea-level changes in the stratigraphic development in the areas around the peninsula. For example, DU 1, DU 2, and DU 3 in the present study correspond respectively to DU 1, DU 2 plus DU 3, and DU 4 in the study of Liu et al. (2007b).

Fig.11 Age-depth plot and sediment accumulation curve for the upper sediments of core QDZ03 (0–8.63 m). 1 sigma ranges of ages are not plotted. A, sea-level curve for the Yellow Sea and adjacent seas (Liu et al., 2004); B, sea-level curve for the Sunda shelf (Hanebuth et al., 2011). DU 1 and DU 2, depositional units.

5.2 Provenance

When plotted on the illite–smectite–(kaolinite+chlorite)ternary diagram (ISKc diagram), clay minerals in the interval 0–8.63 m of core QDZ03 overlap with Yellow River sediments (Fig.12). Illite/smectite ratios are mostly less than 6, values that are consistent with those of Yellow River sediments and different from those of Yangtze River sediments, for which this ratio is generally greater than 8 (Fan et al., 2001). It is therefore reasonable to exclude significant contributions by the Yangtze to the clinoform in the study area. The core interval from 2 to 8.63 m is lower in illite content, and such content in the 0–2 m interval coincides with that for the Yellow River sediments. This pattern points to a growing influence from the Yellow River as its sediment discharge has increased (Saito et al., 2001; Wang et al., 2007).

The distribution pattern of REEs (Fig.10) in the study area suggests that the provenance of sediment in the clinoform is not solely from Yellow River sediments and may also include contributions from proximal sediments. We compared REE data from the sediments of the clinoform (0–8.63 m of core QDZ03) to those from the dominant granite and lamprophyre parent rocks in the Shandong Peninsula (Han et al., 1991, 2010) and from river sediments of the southern peninsula. The REE data of the clinoform are consistent with a mixture of Yellow River sediments and sediments derived locally from these parent rocks and/or the small rivers. The spatial distribution of the clinoform, which thickens at the mouths of Aoshan, Xiaodao, Dingzi, and Rushan bays (Fig.7), also suggests a significant role for local sediment in the development of the clinoform.

Yellow River sediment has been traced in the coastal seas of China by previous studies. Chen et al. (1986) used detrital minerals to show that sediment from the modern Yellow River enters the Yellow Sea through the Bohai Strait, is carried southwest after passing Chengshan Cape, and is deposited as far away as the area of Haiyang City in Shandong Province. Li et al. (2000) found that detrital minerals in the nearshore sediments from Dingzi Bay to Jiaozhou Bay are dominated by potassium feldspar and mica typical of Yellow River detrital minerals. Li et al. (2010) concluded that clay minerals derived from the Yellow River and characterized by the assemblage illite–smectite–(chlorite–kaolinite) are widely distributed in the bottom sediment in the south and north coastal seas of the Shandong Peninsula.

Sources of the clinofrom sediments in the study area can be estimated by a mass balance comparison of the sediment discharges of main rivers flowing from the Southern Shandong Peninsula to the coastal sea with the sediment stored in the clinoform. From the distribution of the thickness of the clinoform (Fig.7), an estimate of the total amount of sediment in the clinoform can be calculated. Within the main body of the clinoform in the area north of Laoshan Head, a total of 2.6×109t of sediment can be accounted for, assuming an average dry bulk density of 1.2 t m-3(Chen et al., 2003). Assuming a relatively constant sediment discharge of 500×103t yr-1for Rushan River to Rushan Bay, 840×103t yr-1for Wulong River to Dingziwan Bay, and 80×103t yr-1for small rivers to Aoshan Bay and Xiaodao Bay (Editorial Board of Annals of Bays in China, 1993), approximately 1.6×109t of sediment has been supplied to the bays by the rivers during the building of the clinoform in a time span of about 11.2 kyr. This estimate of sediment discharges can be considered a maximum if we take it into account the fact that fluvial sediment loads are higher now than before ca. 2000 years BP because of human activities, as evidenced by the cases of the Yangtze Rvier and Yellow River (Saito et al., 2001). Moreover, the estimated sediment discharge (1.6×109t) must have also included sediments transported out of the bays to the open sea by tides, waves, and alongshore currents. The implication is that the total amount of sediment (about 2.6×109t) in the clinoform must have come jointly from the local small rivers and the other sources likely including the Yellow River, as suggested by the analysis above.

Marine currents in the region enable sediment from the Yellow River to be carried through the Bohai Strait and transported around and along the coast of the Shandong Peninsula as far west as the Laoshan Head area (Qin et al., 1989a; Fig.2). Seismic profiles in the study area show that the part of the clinoform that lies above the Holocene MFS generally tilts gently and progrades to the south and east, in accordance with deltaic progradation of sediments from small rivers on the peninsula southward and eastward out into the SYS, coupled with delivery along the coast of sediments derived from the Yellow River.

Fig.12 Clay-mineral ternary diagram (ISKc diagram) showing compositions of sediments from the study area, the Yellow River, and the Yangtze River (data for the Yellow River and Yangtze River from Xu, 1983; Yang, 1988; Yang et al., 2002, 2003a).

5.3 Origin of the Clinoform

As with subaqueous clinoforms elsewhere (Cattaneo et al., 2003; Walsh et al., 2004; Neill and Allison, 2005; Swenson et al., 2005), development of the clinoform in the study area resulted from the action of multiple factors such as sea-level changes, sediment sources, and sedimentary dynamic conditions during the Holocene. Sea level from the last glacial maximum dropped by over 120 m, exposing the continental shelf of the Yellow Sea (Qin et al., 1989b). With the post-glacial sea-level rise, the shoreface moved landward across the shelf, forming aregional diachronous ravinement surface (T1) in the study area in the early Holocene. During this transgression until the middle Holocene sea-level highstand (7–6 cal kyr BP), the Yellow Sea Warm Current and the modern circulation pattern were not established in the Yellow Sea (Liu et al., 1999; Li et al., 2007), and Yellow River sediments discharged into the Bohai Sea were carried by strong tidal currents through the Bohai Strait (Liu et al., 2009) into the North and South Yellow Seas. These sediments accumulated in the study area to form DU 2 in core QDZ03, but the sedimentation rate was relatively low (Fig.11).

After the middle Holocene highstand and the inception of the Yellow Sea Warm Current (Liu et al., 1999; Li et al., 2007) and the modern circulation pattern, coastal currents carried Yellow River sediment to the nearshore area off the southern Shandong Peninsula, forming the main body of the clinoform (DU 1 in core QDZ03). Sediment from the local rivers poured into Rushan, Dingzi, Aoshan, and Xiaodao bays, forming sedimentary fans at their mouths. Because longshore transport of these sediments was blocked at Laoshan Head by vortexes arising from tidal currents (Yuan et al., 2007), the greatest sediment thicknesses are found at the mouths of Aoshan Bay and Xiaodao Bay (Fig.7).

The thickest part of the muddy depositional body is not nearest the coast, because during the Holocene stepped transgression the area next to the coast was not always covered by seawater, and the sediment supply was small. Coastal turbulence also contributed to the resuspension and offshore transport of the muddy sediment. The clinoform northeast of the Shandong Peninsula is likewise not thickest next to the coast (Liu et al., 2007b). From Lao shan Head to the south of Shazi Mouth, there is almost no Holocene sediment (Fig.7), presumably because of rapid sea-level rise, limited sediment supply, and strong tidal action during the Holocene transgression (Yuan et al., 2007). Farther west there is a small area of sediment accumulation off Jiaozhou Bay, part of the ebb-delta lobe of the main tidal channel (Wang et al., 2000). The sedimentary hiatus of about 4000 years between DU 1 and DU 2 in core QDZ03 is construed as the consequence of downlapping of DU 1 of the clinoform onto the MFS.

Previous studies have documented a Holocene clinoform or subaqueous delta northeast of the Shandong Peninsula (Liu et al., 2007b; Yang and Liu, 2007), which is roughly coeval with the clinoform of the study area. Though both are somewhat different in sediment source with the former being composed primarily of Yellow River sediment, the two sedimentary bodies are related in their spatial distribution and Holocene history.

6 Conclusions

We reached the following conclusions from our comprehensive study of about 3000 km of high-resolution shallow seismic profiles and core QDZ03 in the study area:

1) The clinoform off the southern Shandong Peninsula is mainly distributed in the nearshore area to the north of Laoshan Head, with its base generally at 15–40 m bpsl and its thickness ranging from 1 to 15 m. The clinoform gradually thins seaward, but also forms fans of locally greater thickness at the mouths of bays along the coast. The 3-m isopach of the clinoform is an obvious boundary that runs roughly along the present 25-m ocean depth contour. Landward of this contour, the clinoform grows in thickness to a maximum of 22.5 m and becomes steeper, whereas seaward the thickness changes little, ranging between 1 and 3 m.

2) The clinoform in the study area came into being in the early Holocene (about 11.2 cal kyr BP) and can be divided into lower and upper depositional units (DU 2 and DU 1, respectively), bounded by the MFS. DU 2 is construed as a post-glacial transgressive systems tract with a thickness generally less than 3 m due to a relatively low sedimentation rate. DU 1, the main body of the clinoform, lies on the MFS and represents a highstand systems tract dating from middle Holocene time (about 7–6 cal kyr BP) to the present day. In core QDZ03, there is a local sedimentary gap of about 4000 years between DU 2 and DU 1 that is construed as the consequence of downlapping of DU 1 of the clinoform onto the the maximum flooding surface (MFS) of the post-glacial period.

3) Clay mineralogy and REE data show that the provenance of the clinoform sediment in the study area, particularly during the last 1–2 cal kyr BP, is from both the Yellow River and proximal sediments in the coastal zone of the Shandong Peninsula.

4) The clinoform in the study area has a counterpart on the northeast side of the Shandong Peninsula. Both sedimentary bodies have a similar Holocene history.

Acknowledgements

This research was financially supported by the National Natural Science Foundation of China (Nos. 41306063 and 41330964) and by the China Geological Survey (Nos. GZH200900501 and GZH201100203). We are grateful to Profs. Edward A. Laws and Qixiang He for their help in the preparation of this manuscript.

Adamiec, G., and Aitken, M. J., 1998. Dose-rate conversion factors: Update. Ancient TL, 16: 37-50.

Alexander, C. R., DeMaster, D. J., and Nittrouer, C. A., 1991. Sediment accumulation in a modern epicontinental-shelf setting: The Yellow Sea. Marine Geology, 98: 51-72.

Bard, E., Hamelin, B., and Fairbanks, R. G., 1990. U-Th ages obtained by mass spectrometry in corals from Barbados: Sea level during the past 130000 years. Nature, 346: 456-458.

Beardsley, R. C., Limeburner, R., Hu, D., Le, K., Cannon, G. A., and Pashinski, D. J., 1983. Structure of the Changjiang River plume in the East China Sea during June 1980. Proceedings of the International Symposium on Sedimentation on the Continental Shelf with Special Reference to the East China Sea, Hangzhou, China. Vol. 1. China Ocean Press, Beijing, 243-260.

Blanchon, P., and Shaw, J., 1995. Reef drowning during the lastdeglaciation: Evidence for catastrophic sea-level rise and icesheet collapse. Geology, 23: 4-8.

Cattaneo, A., Correggiari, A., Langone, L., and Trincardi, F., 2003. The late-Holocene Gargano subaqueous delta, Adriatic shelf: Sediment pathways and supply fluctuations. Marine Geology, 193: 61-91.

Chen, L., Shen, S., Xu, W., and Li, A., 1986. An approach to the detrital assemblages and their distribution patterns in the sediments of the China Sea. Acta Sedimentologica Sinica, 4 (3): 87-97 (in Chinese with English abstract).

Chen, Z., Saito, Y., Hori, K., Zhao, Y., and Kitamura, A., 2003. Early Holocene mud-ridge formation in the Yangtze offshore, China: A tidal-controlled estuarine pattern and sea-level implications. Marine Geology, 198 (3-4): 245-257.

Cheng, P., Gao, S., Liu, J. P., Guan, C. Z.,Wang, G. Z., and Milliman, J. D., 2001. A preliminary study of the distribution of Holocene deposits in the western North Yellow Sea. Quaternary Sciences, 21: 379 (in Chinese).

Editorial Board of Annals of Bays in China, 1993. Annals of Bays in China (Series 4). China Ocean Press, Beijing, 73-156.

Fairbanks, R. G., 1989. A 17,000-yr glacio-eustatic sea-level record: Influence of glacial melting rates on the Younger Dryas event and deep-ocean circulation. Nature, 342: 637-642.

Fan, D., Yang, Z., Mao, D., and Guo, Z., 2001. Clay minerals geochemistry of the sediments from the Yangtze and Yellow Rivers. Marine Geology & Quaternary Geology, 21 (4): 7-12 (in Chinese with English abstract).

Galbraith, R. F., Roberts, R. G., Laslett, G. M., Yoshida, H., and Olley, J. M., 1999. Optical dating of single and multiple grains of quartz from Jinmium rock shelter, northern Australia. Part I, experimental design and statistical models. Archaeometry, 41: 339-364, DOI: 10.1111/j.1475-4754.1999.tb00987.x.

Han, Z., Li, M., Li, A., Yi, W., and Wang, Q., 2010. REE characteristic and provenance discrimination of sediments from beach north shore of Tianheng Island in Qingdao. Transactions of Oceanology and Limnology, 3: 132-136 (in Chinese with English abstract).

Han, Z., Ni, B., Zhao, G., and Sheng, X., 1991. The elemental geochemistry of Alkali-Granite from Laoshan area. Transactions of Oceanology and Limnology, 1: 16-23 (in Chinese with English abstract).

Hanebuth, T. J. J., Voris, H. K., Yokoyama, Y., Saito, Y., and Okuno, J., 2011. Formation and fate of sedimentary depocentres on Southeast Asia’s Sunda Shelf over the past sea-level cycle and biogeographic implications. Earth-Science Reviews, 104: 92-110.

Jiang, X. J., Lin, X. H., Yao, D., and Guo, W. D., 2011. Enrichment mechanisms of rare earth element in marine hydrogenic ferromanganese crusts. Science China (Earth Science), 54: 197-203, DOI: 10.1007/S11430-010-4070-4.

Li, A., Chen, L., and Li, T., 2000. Mineral assemblage and their provenance and its sedimentary environmental indicative significance of surface sediments in the south Shandong Peninsula. Marine Sciences, 3: 57 (in Chinese).

Li, G., Mu, X., Hu, B., and Peng, W., 2010. Distribution pattern and assemblage feature of clay minerals in surface sediments from the coastal area of Shandong Peninsula. Marine Geology & Quaternary Geology, 30 (4): 62-72 (in Chinese with English abstract).

Li, G., Yang, Z., and Liu, Y., 2005. Genetic Researches on the Submarine Sedimentary Environments of the East China Seas. Science Press, 1-65 (in Chinese).

Li, G., Yue, S., Zhao, D., and Sun, Y., 2004. Rapid deposition and dynamic processes in the modern Yellow River mouth. Marine Geology & Quaternary Geology, 24 (3): 29-35.

Li, T., Jiang, B., Sun, R., Zhang, D., Liu, Z., and Li, Q., 2007. Evolution pattern of warm current system of the East China Sea and the Yellow Sea since the last deglaciation. Quaternary Sciences, 27: 945-954 (in Chinese with English abstract).

Li, X., Li, G., Cui, C., Sun, X., Zhang, W., and Wang, Y., 1994. Relationship between the geological environment and the distribution of warm springs in Shandong Peninsula. Journal of Ocean University of Qingdao, Special issue: 9-15 (in Chinese with English abstract).

Liu, J. P., Milliman, J. D., Gao, S., and Cheng, P., 2004. Holocene development of the Yellow River’s subaqueous delta, North Yellow Sea. Marine Geology, 209: 45-67.

Liu, J. P., Xu, K. H., Li, A. C., Milliman, J. D., Velozzi, D. M., Xiao, S. B., and Yang, Z. S., 2007a. Flux and fate of Yangtze River sediment delivered to the East China Sea. Geomorphology, 85: 208-224.

Liu, J., Li, S., Wang, S., Yang, Z., Ge, Z., and Chang, J. H., 1999. Sea level changes of the Yellow Sea and formation of the Yellow Sea warm current since the last deglaciation. Marine Geology & Quaternary Geology, 19 (1): 13-24 (in Chinese with English abstract).

Liu, J., Saito, Y., Kong, X., Wang, H., and Zhao, L., 2009. Geochemical characteristics of sediment as indicators of post- glacial environment changes off the Shandong Peninsula in the Yellow Sea. Continental Shelf Research, 29: 846-855.

Liu, J., Saito, Y., Kong, X., Wang, H., Wen, C., Yang, Z., and Nakashima, R., 2010. Delta development and channel incision during marine isotope stages 3 and 2 in the western South Yellow Sea. Marine Geology, 278: 54-76.

Liu, J., Saito, Y., Wang, H., Yang, Z., and Nakashima, R., 2007b. Sedimentary evolution of the Holocene subaqueous clinoform off the Shandong Peninsula in the Yellow Sea. Marine Geology, 236: 165-187.

Marsh, R. E., Prestwich, W. V., Rink, W. J., and Brennan, B. J., 2002. Monte Carlo determinations of the beta dose rate to tooth enamel. Radiation Measurement, 35: 609-616, DOI: 10. 1016/S1350-4487(02)00065-3.

Martin, J. M., Zhang, J., Shi, M. C., and Zhou, Q., 1993. Actual flux of the Huanghe (Yellow River) sediment to the western Pacific Ocean. Netherlands Journal of Sea Research, 31: 243-254.

Milliman, J. D., and Meade, R. H., 1983. World-wide delivery of sediment to the oceans. Journal of Geology, 91: 1-21.

Milliman, J. D., Li, F., Zhao, Y. Y., Zheng, T. M., and Limeburner, R., 1986. Suspended matter regime in the Yellow Sea. Progress in Oceanography, 17: 215-227.

Milliman, J. D., Qin, Y., Ren, M. E., and Satio, Y., 1987. Man’s influence on the erosion and transport of sediment by Asian’s rivers. The Yellow River (Huanghe) example. Journal of Geology, 95: 751-762.

Neill, C. F., and Allison, M. A., 2005. Subaqueous deltaic formation on the Atchafalaya Shelf, Louisiana. Marine Geology, 214: 411-430.

Nittrouer, C. A., Kuehl, S. A., Figueiredo, A. G., Allison, M. A., Sommerfield, C. K., Rine, J. M., Faria, L. E. C., and Silveira, O. M., 1996. The geological record preserved by Amazon shelf sedimentation. Continental Shelf Research, 16: 817-841.

Prescott, J. R., and Hutton, J. T., 1994. Cosmic ray contributions to dose rates for luminescence and ESR dating: Large depths and long-term time variations. Radiation Measurement, 23: 497-500, DOI: 10.1016/1350-4487(94)90086-3.

Qin, Y. S., and Li, F., 1986. Study of the influence of sedimentloads discharged from Huanghe River on sedimentation in the Bohai and Yellow Seas. Oceanologia et Limnologia Sinica, 27: 125-135.

Qin, Y. S., Li, F., Xu, S., Milliman, J., and Limeburner, R., 1989a. Suspended matter in the south Yellow Sea. Oceanologia et Limnologia Sinica, 20 (2): 101-112 (in Chinese with English Abstract).

Qin, Y. S., Zhao, Y. Y., Chen, L. R., and Zhao, S. L., 1989b. Geology of the Yellow Sea. China Ocean Press, Beijing, 24-27.

Qin, Y. S., Zhao, Y. Y., Chen, L. R., and Zhao, S. L., 1990. Geology of the Bohai Sea. China Ocean Press, Beijing, 70pp.

Saito, Y., Wei, H., Zhou, A., Sato, Y., and Yokota, S., 2000. Delta progradation and chenier formation in the Huanghe (Yellow River) delta, China. Journal of Asian Earth Science, 18: 489-497.

Saito, Y., Yang, Z., and Hori, K., 2001. The Huanghe (Yellow River) and Changjiang (Yangtze River) deltas: A review on their characteristics, evolution and sediment discharge during the Holocene. Geomorphology, 41: 219-231.

Shi, X., Liu, S., Qiao, S., Liu, Y., Fang, X., Wu, Y., and Zhu, Z., 2010. Deposition features and paleoenvironmental records of the mud deposits in Min-Zhe coastal mud area, East China Sea. Marine Geology & Quaternary Geology, 30 (4): 19-30 (in Chinese with English abstract).

Slingerland, R., Driscoll, N. W., Milliman, J. D., Miller, S. R., and Johnstone, E. A., 2008. Anatomy and growth of a Holocene Clinothem in the Gulf of Papua. Journal of Geophysical Research, 113, FO1S13, DOI: 10.1029/2006JF000628.

Southon, J., Kashgarian, M., Fontugne, M., Metivier, B., and Yim, W. W. S., 2002. Marine reservoir corrections for the Indian Ocean and southeast Asia. Radiocarbon, 44: 167-180.

Stuiver, M., Reimer, P. J., and Reimer, R. W., 2005. CALIB 5.0.1 [WWW program and documentation] (http://calib.qub.ac.uk/). Su, J. L., and Yuan, Y. L., 2004. China Offshore Hydrology. China Ocean Press, Beijing, 174-206.

Swenson J. B., Paola, C., Pratson, L., Voller, V. R., and Murray, A. B., 2005. Fluvial and marine controls on combined subaerial and subaqueous delta progradation: Morphodynamic modeling of compound-clinoform development. Journal of Geophysical Research, 110, F0213, DOI: 10.1029/2004JF000265.

Ta, T. K. O., Nguyen, V. L., Tateishi, M., Kobayashi, I., Tanabe, S., and Saito, Y., 2002a. Holocene delta evolution and sediment discharge of the Mekong River, southern Vietnam. Quaternary Science Reviews, 21: 1807-1819.

Ta, T. K. O., Nguyen, V. L., Tateishi, M., Kobayashi, I., Tanabe, S., Saito, Y., and Nakamura, T., 2002b. Sediment facies and Late Holocene progradation of the Mekong River Delta in Bentre Province, southern Vietnam: An example of evolution from a tide-dominated to a tide- and wave-dominated delta. Sedimentary Geology, 152: 313-325.

Tornqvist, T. E., Bick, S. J., and Gonzalez, J. L., 2004. Tracking the sea-level signature of the 8.2 ka cooling event: New constraints from the Mississippi Delta. Geophysical Research Letters, 31, L23309, DOI: 10.1029/2004GL021429.

Walsh, J. P., Nittrouer, C. A., Palinkas, C. M., Ogston, A. S., Sternberg, R. W., and Brunskill, G. J., 2004. Clinoform mechanics in the Gulf of Papua, New Guinea. Continental Shelf Research, 24: 2487-2510.

Wang, H. J., Yang, Z. S., Saito, Y., Liu, J. P., Sun, X. X., and Wang, Y., 2007. Stepwise decrease of the Huanghe (Yellow River) sediment load (1950–2005): Impacts of climate change and human activities. Global and Planetary Change, 57: 331-354.

Wang, H. L., and Li, G. S., 2009. Numerical simulation on seasonal transportation of suspended sediment from Huanghe (Yellow) River to Bohai Sea. Oceanologia et Limnologia Sinica, 40 (2): 129-137 (in Chinese with English Abstract).

Wang, P., Min, Q., and Bian, Y., 1985. Distributions of foraminifera and ostracoda in bottom sediments of the northwestern part of the South Huanghai (Yellow) Sea and its geological significance. In: Marine Micropaleontology of China. Wang, P., et al., eds., China Ocean Press, Beijing, 93-114.

Wang, Y., Gao, S., and Jia, J., 2000. Sediment distribution and transport patterns in Jiaozhou Bay and adjoining areas. Acta Geographica Sinica, 55: 449-458 (in Chinese with English abstract).

Weaver, A. J., Saenke, O. A., Clark, P. U., and Mitrovica, J. X., 2003. Meltwater pulse 1A from Antarctica as a trigger of the Bølling-Allerød warm interval. Science, 299: 1709-1713.

Xu, D. Y., 1983. Mud sedimentation on the East China Sea shelf. In: Proceeding of International Symposium on Sedimentation on the Continental Shelf with Special Reference to the East China Sea. 12–16 April 1983. China Ocean Press, Beijing.

Yang, S. Y., and Li, C. X., 1999. REE geochemistry and tracing application in the Yangtze River and the Yellow River sediments. Geochimica, 28: 374-380 (in Chinese with English abstract).

Yang, S. Y., Jung, H. S., Choi, M. S., and Li, C. X., 2002. The rare earth element compositions of the Changjiang (Yangtze) and Huanghe (Yellow) river sediment. Earth and Planetary Science Letters, 201: 407-419.

Yang, S. Y., Jung, H. S., Lim, D. I., and Li, C. X., 2003a. A review on the provenance discrimination in the Yellow Sea. Earth-Science Reviews, 63: 93-120.

Yang, S. Y., Li, C. X., Jung, H. S., and Choi, M. S., 2003b. Constraints and tracer significance rethinking of the REE in the Yellow River sediments. Advance in Nature Sciences, 13: 365-371 (in Chinese).

Yang, S. Y., Li, C. X., Lee, C. B., and Na, T. K., 2003c. REE geochemistry of suspended sediments from the rivers around the Yellow Sea and provenance indicators. Chinese Science Bulletin, 48: 1135-1139.

Yang, Z. S., 1988. Mineralogical assemblages and chemical characteristics of clays from sediments of the Huanghe, Changjiang, Zhujiang rivers and their relationship to the climate environment in their sediment source areas. Oceanologia et Limnologia Sinica, 19: 336-346 (in Chinese with English Abstract).

Yang, Z. S., and Liu, J. P., 2007. A unique Yellow River-derived distal subaqueous delta in the Yellow Sea. Marine Geology, 240: 169-176.

Yao, Y., Xia, B., Feng, Z., Wang, L., and Xu, X., 2005. Tectonic evolution of the South Yellow Sea since the Paleozoic. Petroleum Geology & Experiment, 27 (2): 124-128 (in Chinese with English abstract).

Ying, Z. F., 1999. On the coast current and its deposit along western coast in Guangdong. Acta Scientiarum Naturalium University Sunyatseni, 38 (3): 85-89 (in Chinese with English abstract).

Yuan, H., Li, S., Dong, H., and Li, C., 2007. On the offshore depositional system in Qingdao. Marine Geology & Quaternary Geology, 27 (Suppl.): 74-82 (in Chinese).

Zhang, R., 1984. Land-forming history of the Huanghe River Delta and coastal plain of northern Jiangsu. Acta Geographica Sinica, 39: 173-184 (in Chinese with English abstract).

(Edited by Xie Jun)

(Received November 20, 2012; revised December 11, 2012; accepted December 20, 2013)

© Ocean University of China, Science Press and Springer-Verlag Berlin Heidelberg 2014

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