Magnetostratigraphy of a long Quaternary sediment core in the South
Yellow Sea
Jianxing Liu a, b, c, d
, Qingsong Liu c, d, *
, Xunhua Zhang b, **
, Jian Liu b, d
, Zhiqiang Wu b
,
Xi Mei b, d
, Xuefa Shi a, d
, Quanhong Zhao e
a
Key Laboratory of Marine Sedimentology and Environmental Geology, First Institute of Oceanography, State Oceanic Administration, Qingdao, 266061,
People’s Republic of China
b
Qingdao Institute of Marine Geology, Qingdao, 266071, People’s Republic of China
c
State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing, 100029, People’s Republic of
China
d
Laboratory for Marine Geology, Qingdao National Laboratory for Marine Science and Technology, Qingdao, 266061, People’s Republic of China
e
State Key Laboratory of Marine Geology, Tongji University, Shanghai, 200092, People’s Republic of China
a r t i c l e i n f o
Article history:
Received 4 January 2016
Received in revised form
9 May 2016
Accepted 18 May 2016
Available online 27 May 2016
Keywords:
Continental shelf
The South Yellow Sea
Core CSDP-1
Greigite
Magnetostratigraphy
Quaternary
Sedimentary evolution
a b s t r a c t
Continental shelves serve as a bridge between the continent and ocean and sediments in this region are
sensitive to land-sea interaction, sea-level variation and local subsidence. In this study, we present a
comprehensive magnetic study of the longest sediment core (CSDP-1, 300.1 m) recovered from the South
Yellow Sea. The major magnetic minerals in the studied sediments are magnetite, hematite and greigite.
Greigite records a chemical remanent magnetization, which can be removed effectively by thermal
demagnetization. The magnetostratigraphy defined in this study contains the Matuyama-Brunhes
boundary (M/B, 781 ka) at ~73.68 m, which is consistent with results from adjacent cores. The base of
the Quaternary (~2.6 Ma) in the Yellow Sea is recovered for the first time at a depth of 227.16 m. The basal
age of the core is estimated to be ~3.50 Ma. It indicates that the first transgression of the Yellow Sea
occurred no later than ~1.7 Ma. Succeeding large amplitude regressions occurred in some cold periods
such as during MIS 20, MIS 18, and MIS 10. Our results provide the first chronology that brackets the
entire Quaternary and we reconstruct the sedimentary evolution of the Yellow Sea with robust age
constraints, which provides an important framework for further paleoenvironmental and tectonic
studies.
© 2016 Elsevier Ltd. All rights reserved.
1. Introduction
Continental shelves are among the most active regions for landsea
interaction in the Earth System (McMillan, 2002), and are
typically important for economic production and for human recreational
activities (Eckert, 1979). Sediments from the continental
shelf are highly sensitive to sea-level fluctuations, climate changes,
and local subsidence (Troiani et al., 2011). Fluvial material that is
transported to this region can either accumulate or be retransported
to the ocean depending on the accommodation space
that is controlled by both sea-level and tectonic subsidence (Gerber
et al., 2010; Yao et al., 2014). Sedimentation rates on the continental
shelf are usually much higher than those in the deep sea, which
makes shelf sediments suitable for high-resolution studies (e.g., Liu
et al., 2004).
The Yellow Sea (Fig. 1) is a typical continental shelf with
extensive area in the east of China. Deposited under relatively
shallow water depths (44 m on average), sediment sequences in
this region are highly sensitive to sea-level fluctuations as a result
of Quaternary glacial-interglacial cycles as well as local tectonics
(Qin et al., 1989; Yang et al., 1996). In addition, it is an important
sediment reservoir for two globally large rivers, the Yellow River
and Yangtze (Changjiang) River, and other small rivers on the
Korean Peninsula (Yang et al., 2003). Therefore, sediments from the
* Corresponding author. State Key Laboratory of Lithospheric Evolution, Institute
of Geology and Geophysics, Chinese Academy of Sciences, Beijing, 100029, People’s
Republic of China.
** Corresponding author. Qingdao Institute of Marine Geology, Qingdao, 266071,
People’s Republic of China
E-mail addresses: qsliu@mail.iggcas.ac.cn (Q. Liu), xunhuazh@vip.sina.com
(X. Zhang).
Contents lists available at ScienceDirect
Quaternary Science Reviews
journal homepage: www.elsevier.com/locate/quascirev
http://dx.doi.org/10.1016/j.quascirev.2016.05.025
0277-3791/© 2016 Elsevier Ltd. All rights reserved.
Quaternary Science Reviews 144 (2016) 1e15
Yellow Sea will contain important paleoenvironmental information
that is relevant to studying sea-level and tectonic variations, source
to sink processes, and East Asian monsoon evolution (Liu et al.,
2014a). Nevertheless, most previous studies have focused on
shorter time-scales over the last interglacial, such as evolution of
the Yellow River subaqueous delta during the Holocene (Yang and
Liu, 2007), marine isotope stage (MIS) 3 to MIS 2 (Liu et al., 2010),
and the paleoenvironmental history of the central South Yellow Sea
(SYS) mud area since the early Holocene (Xiang et al., 2008). Studies
on longer time-scales are relatively scarce due to the lack of reliable
chronological frameworks.
For long-core sediment sequences, magnetostratigraphy is one
of the most efficient ways to establish a reliable first-order chronological
framework (e.g., Lanci et al., 2004; Tauxe et al., 2012). So
far, magnetostratigraphic studies have been reported for four long
cores (QC2, EY02-2, NHH01, and DLC70-3) (Fig. 1b; Table 1) from
the Yellow Sea. The 108.83-m QC2 core contains the MatuyamaBrunhes
boundary (M/B: ~781 ka B.P.) at 79.95 m (or ~63 m after
removing the ~17-m-thick Holocene sand ridge) and a basal age of
~1.8 Ma (Zhou and Ge, 1990). The M/B boundary of core EY02-2,
which has a total length of 70 m, was identified at 63.29 m, and
its basal age was extrapolated to ~0.9 Ma (Ge et al., 2006). Recent
studies have located the M/B boundary in the 125.64-m NHH01
core at 68.64 m with a linearly extrapolated basal age of ~1.06 Ma
(Liu et al., 2014a), and the M/B boundary of the 71.2-m DLC70-3
core at 59.08 m (Mei et al., 2016).
Quaternary climate is characterized by glacial and interglacial
alternations that have resulted in large-scale sea-level oscillations
that influenced the architecture of continental shelf sediments (Yao
et al., 2014). Numerous sediment cores from the Yellow Sea and its
surrounding coastal plains have revealed that sedimentation in this
region is characterized by alternation of marine and lacustrine/
terrestrial deposits (Qin et al., 1989). In core QC2 (Fig. 1b), seven
marine transgressions divide the recorded paleogeographic evolution
into four periods since the Olduvai subchron (Yang, 1993).
Comprehensive analyses of seismic profiles and sediment cores
from the western SYS reveal that sedimentary environmental
changes since MIS 5 were strongly controlled by sea-level fluctuations,
with most of the preserved sediments deposited in MIS 5,
MIS 3, and MIS 1 and major erosion events in MIS 4 and MIS 2 (Liu
et al., 2010). Using foraminifera abundances, eight transgression
layers have been identified in core DLC70-3 (Fig. 1b) (Mei et al.,
2016). In short, great progress has been made in understanding
the sedimentary evolution of the Yellow Sea since the middle
Pleistocene, especially over the late Pleistocene. However, the
sedimentary evolution history throughout the Quaternary is still
not well understood.
In this study, we investigated a 300.1-m-long sediment core
from the SYS. Detailed rock magnetic and high-resolution paleomagnetic
investigations are presented along with a robust chronology
that brackets the entire Quaternary for the SYS. We use this
paleoenvironmental reconstruction to outline the sedimentary
history of the SYS.
2. General setting
The Bohai Sea, Yellow Sea, and the East China Sea together
constitute a marginal sea bounded by China, Korea, and southern
Fig. 1. (a). Schematic map of the marginal sea distribution in the northwest Pacific Ocean (BS: the Bohai Sea; YS: the Yellow Sea; ECS: the East China Sea; SCS: the South China Sea).
(b). Map of the Yellow Sea and adjacent areas, with locations of existing long sediment cores, and the cores QC2, EY02-2, NHH01, DLC70-3 are from Zhou and Ge (1990), Ge et al.
(2006), Liu et al. (2014a), and Mei et al. (2016), respectively.
Table 1
Detailed information on long sediment cores from the South Yellow Sea.
Core ID Water depth/m Drilling depth/m Recovery rate/% Location longitude/latitude Drilling time
QC2 49.05 108.83 90.4 122
160
E/34
180
N May, 1984
EY02-2 79.00 70.00 86.5 123
300
E/34
300
N February, 2001
NHH01 73.00 125.64 91.0 123
130
E/35
130
N June, 2009
DLC70-3 73.00 71.20 93.0 123
330
E/36
380
N September, 2009
CSDP-1 52.50 300.10 80.0 122
220
E/34
180
N June, 2013
J. Liu et al. / Quaternary Science Reviews 144 (2016) 1e152
Japan, and is separated from the western North Pacific Ocean by the
Ryukyu Island arc and Taiwan (Fig. 1a, Liu et al., 2010). More than
70% of this marginal sea is occupied by a broad continental shelf
that is the seventh largest in the world with water depths of less
than ~150 m. The Yellow Sea is a typical semi-enclosed (open to the
southeast) epicontinental sea that rests on a flat and tectonically
stable seafloor, with average and maximum water depths of ~44 m
and 140 m, respectively (Yang et al., 2003; Liu et al., 2014a). It is
located between the Chinese mainland and Korean Peninsula, it
stretches 870 km from north to south and 556 km from east to
west, and it has an area of ~380,000 km2
(He, 2006).
The link between the Chengshan Cape of Shandong Peninsula
and Changsangot of the Korean Peninsula divides the Yellow Sea
into the North Yellow Sea (NYS) and SYS. The NYS is separated from
the Bohai Sea at its western extremity by Liaotieshan Cape of
Liaodong Peninsula and Penglai Cape of Shandong Peninsula, while
the SYS is bounded in the south with the East China Sea by an
arbitrary line that connects the Yangtze River Estuary and Cheju
Island (Fig. 1b, Qin et al., 1989). Water depths in the NYS are
generally less than 60 m with an average of 38 m, and the SYS is
mostly less than 100 m deep with an average water depth of 46 m.
The SYS has a maximum width of ~700 km and deepens toward the
Yellow Sea Trough, which is defined by the 80-m isobath and extends
NW-SE (Liu et al., 2010).
The SYS basin has undergone extensive subsidence during the
Cenozoic so that it contains an enormous thickness of sediments,
with Quaternary sediments having thickness of up to 300 m (Qin
et al., 1989). Mineralogical and geochemical studies of the surface
sediments indicate that the major source of sediments to the
western and middle part of the Yellow Sea are the Yellow River and
Yangtze River, while Korean rivers dominate in the east (Lan et al.,
2005).
3. Material and methods
3.1. Core description
The studied core CSDP-1 (122220 E, 34180 N, Fig. 1b; Table 1)
was recovered from the western SYS at a water depth of ~52.5 m.
Core CSDP-1 is the first borehole of the “China Continental Shelf
Drilling Program” that started in 2011 and is also the longest
sediment core recovered from the SYS so far. The core was recovered
using rotary drilling method in June, 2013, with a total drilling
length of 300.1 m and an average recovery rate of 80%.
The recovered core was split equally into two halves, which
were subsequently photographed and described. The archive half
was preserved in a 4 C repository and the working half was
sampled. According to core observations and descriptions, silt and
clayey silt dominate the sediment sequence, with occasional
interbedded medium-coarse sand/gravel layers. Marine sediments
are only identified in the upper half of the core, mainly in the upper
85 m of the recovered section.
3.2. Rock magnetic experiments
Hysteresis loops were measured on 567 samples at ~0.5 m intervals
using a Princeton Measurements Corporation vibrating
sample magnetometer (VSM, MicroMag 3900). The magnetic field
was cycled between þ1.5 T and À1.5 T, with a field increment of
5 mT and an averaging time of 300e500 ms. Saturation magnetization
(Ms), saturation remanence (Mrs), and coercivity (Bc) were
determined after paramagnetic slope correction. Specimens were
then demagnetized in alternating fields up to 1.5 T, and an
isothermal remanent magnetization (IRM) was imparted stepwise
from 0 to 1.5 T. Finally, backfields were applied stepwise up to
À1.0 T to the magnetically saturated specimens to obtain the
coercivity of remanence (Bcr). For representative samples, firstorder
reversal curves (FORCs) were also measured using the VSM
3900, with 100 FORCs measured at fields up to 1.5 T and an averaging
time of 500 ms. FORC diagrams (Pike et al., 1999) were then
processed using the FORCinel software of Harrison and Feinberg
(2008).
Temperature-dependent curves of magnetic susceptibility (c-T,
mass specific) were measured for 16 representative samples using a
Kappabridge MFK1-FA system with a CS-3 high-temperature
furnace (Agico Ltd., Brno) from room temperature to 700 C in an
argon atmosphere (Hrouda,1994). At parallel depths, cubic samples
were subjected to analysis of thermal demagnetization of threecomponent
IRMs following Lowrie (1990). The samples were
magnetized in successively smaller fields along three mutually
orthogonal axes using a 2-G Enterprises Pulse Magnetizer (2G660),
with a 2.5 T induction along the z-axis (high-coercivity fraction),
0.5 T along the y-axis (medium-coercivity fraction), and 0.05 T
along the x-axis (low-coercivity fraction), respectively. Then samples
were subjected to progressive thermal demagnetization up to
690 C at 20e50 C intervals.
Low-temperature measurements were conducted on splits of
these representative samples using a Quantum Design Magnetic
Property Measurement System (MPMS-5XL). Samples were first
cooled from 300 K to 20 K in zero field, then an IRM was imparted
in a field of 2.5 T. After removal of the field, the IRM was thermally
demagnetized by sweeping the temperature from 20 K to 300 K at
5 K/min.
Anistropy of magnetic susceptibility (AMS) for all paleomagnetic
samples was measured using a KLY-4s Kappabridge instrument
before the samples were subjected to demagnetization. AMS can be
described in terms of an ellipsoid with three mutually orthogonal
susceptibility axes that are designated the maximum (K1), intermediate
(K2), and minimum (K3) axes. These quantities can be
combined variously to describe the shape of the ellipsoid and the
features of the corresponding magnetic fabric (Hrouda, 1982). The
major AMS parameters examined in this study are the lineation (L,
defined as L ¼ K1/K2) and foliation (F, defined as F ¼ K2/K3).
3.3. Paleomagnetic measurements
Stepwise alternating field (AF) demagnetization of the natural
remanent magnetization (NRM) was performed on 2153 cubic
samples at a sampling interval of ~0.1 m using a 2-G Enterprises
Model 760-R cryogenic magnetometer (2G760) installed in a
magnetically shielded (<300 nT) space. AF demagnetization steps
of 5e10 mT were used up to a maximum AF of 100 mT.
Thermal demagnetization of the NRM and remanence measurements
were carried out on 523 samples collected with
2 Â 2 Â 2 cm nonmagnetic quartzose boxes at ~0.5 m sampling
intervals also in the magnetically shielded laboratory. The samples
were heated to a maximum temperature of 690 C at 20e50 C
intervals in a self-made thermal demagnetizer (PGL100) with much
higher thermal stability and lower residual field (AC current field is
~800 nT) than similar commercially available instruments (e.g., AC
current field of TD48 is ~80 mT) (Zheng et al., 2010). Remanence was
measured in a 2-G Enterprises Rapid System in which the minimum
background measurements (~1 Â 10À12
Am2
) are much smaller
than for conventional magnetometers. All magnetic measurements
were accomplished in the Paleomagnetism and Geochronology
Laboratory, Institute of Geology and Geophysics, Chinese Academy
of Sciences.
J. Liu et al. / Quaternary Science Reviews 144 (2016) 1e15 3
4. Results
4.1. Rock magnetism
Hysteresis loops, IRM acquisition curves and FORC diagrams
were used to assess the magnetic mineralogy of the studied sediments
(Fig. 2). For the majority of samples, their hysteresis loops are
closed prior to 0.5 T (e.g., Fig. 2aec) and the corresponding IRM
acquisition curves are almost saturated at ~0.3 T (e.g., Fig. 2eeg),
which indicates that low-coercivity minerals dominate the magnetic
mineral assemblages. Some samples have relatively higher Bc
values (25e56 mT) with Bcr/Bc ~1.5 and Mrs/Ms ~0.5 (e.g., Fig. 2a, b),
which is indicative of typical single-domain (SD) behavior. In
addition, the corresponding FORC diagrams of these samples are
characterized by closed concentric contours about a central peak, a
negative region in the lower left-hand part of the diagram, a peak
value of 60e70 mT for Bc, and an evident spread in the Bu direction
(e.g., Fig. 2i, j). All of these lines of evidence indicate the presence of
authigenic SD greigite (e.g., Roberts, 1995; Sagnotti et al., 2010;
Roberts et al., 2006, 2011; Liu et al., 2014b). In contrast, FORC diagrams
for other samples are typical of low-coercivity SD-PSD
(pseudo-single domain) and/or PSD-MD (multi-domain) magnetic
minerals (Roberts et al., 2000, 2014; Muxworthy and Dunlop, 2002;
Qin et al., 2008) (e.g., Fig. 2k). For some samples with low
magnetization, hysteresis loops are wasp-waisted and are still not
saturated at high fields (e.g., Fig. 2d). Together with unsaturated
IRM acquisition curves (e.g., Fig. 2h) up to 1.5 T, this suggests the
coexistence of high- (e.g., hematite) and low-coercivity (e.g.,
magnetite) phases (Roberts et al., 1995). FORC diagrams for this
kind of sample are noisy and only contain a weak signal associated
with the low-coercivity component (e.g., Fig. 2l).
Temperature-dependent magnetic susceptibility (c-T) measurements
can be further used to determine magnetic minerals in
samples (Hrouda, 1994; Dunlop and €Ozdemir, 1997). c-T curves for
most samples are characterized by a marked drop in c at about
585 C (e.g., Fig. 3aec), the Curie point of magnetite. Other c-T
curves contain a major drop at about 680 C (e.g., Fig. 3d), which is
characteristic of hematite. In addition, increased c on heating from
about 350 C in some samples (e.g., Fig. 3aec) likely results from
neoformation of strong ferrimagnetic minerals via alteration of
iron-containing silicates/clays (e.g., Hirt and Gehring, 1991; Hirt
et al., 1993) and/or paramagnetic iron sulfides such as pyrite (e.g.,
Passier et al., 2001) which usually coexists with greigite (e.g.,
Roberts et al., 2011). The cooling cycles of c-T curves for all samples
Fig. 2. Rock magnetic results for representative samples, including (aed) hysteresis loops after paramagnetic slope correction, (eeh) IRM acquisition curves, and (iel) first-order
reversal curve (FORC) diagrams which are obtained by measuring 100 FORCs in an applied field up to 1.5 T and an averaging time of 500 ms, and then calculated with a smoothing
factor (SF) of 5. The hysteresis loops and IRM curves are obtained in an applied field up to 1.5 T, with a field increment of 5 mT and an averaging time of 300 ms, and the unit for Bc
and Bcr in (aed) is mT.
J. Liu et al. / Quaternary Science Reviews 144 (2016) 1e154
increase when cooled below 585 C (Fig. 3aed), which indicates
that magnetite was produced during the heating/cooling process.
Thermal demagnetization curves for three-component IRMs
indicate that low-, medium- and high-coercivity minerals all
contribute to the remanence (Fig. 3eeh). For some samples, the
maximum unblocking temperature of the dominant IRM carrier is
in the range of 300e400 C (e.g., Fig. 4eef), which along with
hysteresis parameters (Fig. 3aeb) and FORC diagrams (Fig. 3i, j)
undoubtedly indicate the presence of greigite (Chang et al., 2008;
Sagnotti et al., 2010; Roberts et al., 2011). The IRM can be fully
demagnetized at around 600 C, which indicates the coexistence of
greigite and magnetite. For the other samples, the dominant IRM of
some samples is carried by the low- and medium-coercivity components
that evidently unblock at ~600 C (e.g., Fig. 3g), which
indicates that magnetite is the dominant magnetic mineral. For the
hard-component, the IRM is gradually demagnetized up to ~680 C
(e.g., Fig. 3h), which indicates that contributions from hematite are
significant.
Low-temperature measurements on some representative samples
contain an evident Verwey transition at 120 K (Verwey, 1939),
which indicates the presence of stoichiometric magnetite (e.g.,
Fig. 3j, k). In contrast, the Verwey transition is highly smeared (e.g.,
Fig. 3i, l) in other samples, which indicates the minor presence of
magnetite. These results are well consistent with those of the
Fig. 3. Rock magnetic results for representative samples, including (aed) temperature-dependent curves of magnetic susceptibility (c-T), (eeh) thermal demagnetization of threecomponent
IRMs (Lowrie, 1990) produced by magnetizing specimens in 2.5, 0.5 and 0.05 T along the z-, y-, and x-axes, respectively, and (iel) low-temperature curves for an IRM
(mass specific) acquired at 20 K by applying a 2.5-T field and then observed at 5 K intervals to 300 K during zero-field warming (blue lines), with first-order derivative of magnetizations
with respect to temperature (dM/dT) ploted on the right y-axes (magenta lines). The red and blue lines in (aed) denote heating and cooling curves, respectively. (For
interpretation of the references to color in this figure caption, the reader is referred to the web version of this article.)
J. Liu et al. / Quaternary Science Reviews 144 (2016) 1e15 5
thermal demagnetization of three-component IRMs (Fig. 3eeh).
In the AMS results, L is usually smaller than F, which indicates
that the AMS ellipsoid is oblate (Fig. 4b, d). More than 60% of the
samples treated with AF demagnetization (Fig. 4a) and 80% of the
samples treated with thermal demagnetization (Fig. 4c) have subvertical
minimum susceptibility axes (K3) that are almost perpendicular
to the bedding plane, while the inclinations of the
maximum axes (K1) are relatively shallow (<20). These results
denote an original sedimentary magnetic fabric that has remained
unperturbed since deposition for most samples.
4.2. Paleomagnetism
The characteristic remanent magnetization (ChRM) of the
magnetite-dominant samples can be identified mostly between
20 mT and 60 mT for AF demagnetization (e.g., Fig. 5a, d) or between
450 C and 600 C for thermal demagnetization (e.g., Fig. 5f,
h), respectively. For some greigite-bearing samples, the ChRM
could also be isolated using AF demagnetization; at least 80% of
NRM was removed below 80 mT (e.g., Fig. 5b-c), whereas the others
acquired a spurious gyromagnetic remanent magnetization (GRM)
when demagnetized to high fields (Fig. 6a, c) as has been reported
widely to be more typical of greigite (e.g., Snowball, 1997; Roberts
et al., 2011).
Demagnetization results were evaluated using orthogonal diagrams
(Zijderveld, 1967), and each ChRM direction was computed
by a “least-squares fitting” technique (Kirschvink, 1980) with at
least four successive data points that define a linear trend that is
directed clearly to the origin of the orthogonal plot. Samples with
maximum angular deviation (MAD) > 15 or/and K3
inclinations < 50 were excluded. The core was not oriented, so
only magnetic inclination (Inc) data were subsequently used to
construct a magnetostratigraphy.
Paleomagnetic inclinations calculated from the AFD or thermal
demagnetization data prior to 400 C are usually different from
those computed from the thermal demagnetization data above
400 C in the greigite-dominated samples as we also encountered
in core NHH01 (Fig. 1b, Liu et al., 2014a). For example, the paleomagnetic
inclinations for some samples have the same direction
but distinctly different values (e.g., Fig. 5e) while in other samples,
Fig. 4. Anisotropy of magnetic susceptibility (AMS) data for samples subjected to (a,b) AF and (c,d) thermal demagnetization, respectively. (a,c) Projection of axes of maximum (K1)
and minimum (K3) susceptibility. (b,d) Magnetic lineation (L) versus magnetic foliation (F).
J. Liu et al. / Quaternary Science Reviews 144 (2016) 1e156
Fig. 5. Orthogonal vector plots (left) and remanence decay curves (right) for (aed) stepwise AF and (eeh) thermal demagnetization data for representative samples. Solid (open)
circles represent projections onto the horizontal (vertical) plane. The green dashed lines in the remanence decay curves in (aed) denote the median destructive field (MDF). The
dashed lines in (e) represent fits to two separate components. The core was not oriented, so no declinations are reported. “Step” indicates the demagnetization range over which the
linear component fitting was conducted for each sample. (For interpretation of the references to color in this figure caption, the reader is referred to the web version of this article.)
J. Liu et al. / Quaternary Science Reviews 144 (2016) 1e15 7
the inclinations changed sign when heated to ~360 C (e.g., Fig. 6b,
d). Although the Curie temperature of greigite remains unknown
but must exceed 350 C, it will be almost completely demagnetized
prior to 400 C (Roberts et al., 2011). Therefore, the NRM that remains
above this temperature interval will be carried only by iron
oxides (mainly magnetite).
Three magnetozones are recognized in the core (Fig. 7b): two
with normal polarity, N1 (0e73.68 m) identified from thermal
demagnetization data above 400 C or (0e82.80 m) by both AF data
and thermal demagnetization data below 400 C, N2
(227.16e300.10 m); and one with reverse polarity, R1
(73.68e227.16 m) or (82.80e227.16 m) as determined from both
demagnetization types. Moreover, eight negative inclination
anomalies are recognized within N1 and N2, labeled as r1
(13.47e13.99 m), r2 (18.86e19.25 m), r3 (21.76e23.05 m), r4
(25.75e26.08 m), r5 (41.16e43.28 m), which is not observed in
thermal demagnetization data above 400 C, r6 (57.17e57.96 m), r7
(255.01e257.78 m), and r8 (269.45e272.38 m). R1 is also punctuated
by four relatively thin intervals with positive inclinations,
which are labeled as n1 (93.08e93.50 m), n2 (111.90e114.26 m), n3
(140.12e140.86 m), and n4 (154.93e160.53 m) in Fig. 7b.
5. Discussion
5.1. Remanence recording mechanism in greigite
Paleoenvironments (e.g., redox conditions) on the Chinese
continental shelf change at different time scales due to the combined
effects of organic input, sea-level changes and local tectonics
(Liu et al., 2014a; Mei et al., 2016). Early diagenetic reactions are the
most familiar means by which postdepositional alteration of
detrital magnetic minerals (magnetite and hematite) occurs in
marine sediments (e.g., Larrasoa~na et al., 2007; Fu et al., 2008). It
has been reported that non-steady state diagenesis is a major factor
in sediments from the continental shelf where large-amplitude
sea-level variations give rise to variable changes in organic carbon
content, sulphate content and pore water content with time
(Roberts, 2015). If such sediments have been subjected to continuous
sulphate reduction, detrital magnetic minerals such as
magnetite and hematite will dissolve (e.g., Canfield and Berner,
1987). Oscillation between oxic and sulphidic diagenetic conditions
is likely in such environments. Accordingly, magnetic mineral
assemblages in the sediments of this region are complex and
include both primary iron oxides such as magnetite and hematite,
and secondary iron sulfides such as greigite (e.g., Oda and Torii,
2004; Nilsson et al., 2013; Liu et al., 2014b).
Authigenic greigite forms in sulfate-reducing sedimentary environments
as an intermediate phase during pyritization reactions
(e.g., Berner, 1984; Benning et al., 2000). Some studies suggest that
pyritization reactions occur during early diagenesis at shallow
burial depths and that the persistence of greigite in geological records
can provide an approximately syn-depositional paleomagnetic
record (e.g., Roberts and Turner, 1993; Reynolds et al., 1999).
In contrast, other studies have demonstrated that delayed formation
of greigite gives rise to late sedimentary remagnetizations
(Florindo and Sagnotti,1995; Jiang et al., 2001; Roberts and Weaver,
2005). Greigite can authigenically form at various depths below the
sediment-water interface during diagenesis, which makes it difficult
to ascertain the time at which the chemical remanent
magnetization (CRM) was acquired (e.g., Florindo and Sagnotti,
1995; Liu et al., 2004; Roberts et al., 2005; Roberts and Weaver,
2005; Sagnotti et al., 2005, 2010; Roberts, 2015). The timing of
greigite formation can vary enough to give contradictory polarities
for different generations of greigite even within a single stratigraphic
horizon (Jiang et al., 2001). A second possible explanation
proposed for the inconsistent remanences is self-reversal of greigite
(Horng et al., 1998). However, the physical mechanism for this
hypothetical process has remained poorly understood (Florindo
and Sagnotti, 1995).
Unlike self-reversal which usually occurs in titanomagnetites
and results from exchange interactions within a particle (e.g.,
O’Reilly and Banerjee, 1966; Schult, 1968), Liu et al. (2014a) proposed
that vortex magnetic states in magnetite particles can be
expected to realign the magnetization of nearby SD greigite particles
and that it could give rise to recording of an opposite polarity
direction with respect to the ambient field. Subsequently, the larger
magnetite particles might be partially or completely dissolved
depending on the supply of sulphide (H2S or HSÀ
) that produced by
sulphate reduction (Canfield and Berner, 1987; Roberts, 2015),
leaving greigite as the major remanence carrier. This model gives
rise to “magnetic interaction-induced contradictory magnetization”
and predicts that authigenic greigite of either early or late
diagenetic origin can record various polarities according to the
spatial arrangement of interactions among SD greigite and larger
magnetite particles (Liu et al., 2014a).
Greigite in the studied core is authigenic according to the rock
magnetic results (see Fig. 2i, j). Greigite-bearing layers (grey bars)
(Fig. 7e) often record remanence directions with both the same and
opposite polaries with respect to coexisting magnetite or/and hematite
(see Figs. 5ee7c), which supports the “magnetic interactioninduced
contradictory magnetization” mechanism. Nevertheless,
thermal demagnetization can efficiently remove the effects of
greigite on paleomagnetic directions, so that only magnetite and
hematite signals represent the primary depositional remanent
magnetization.
5.2. Magnetostratigraphy and paleoenvironments of the studied
core
According to the above analyses, the M/B boundary in the
studied core is confidently identified at 73.68 m by combining both
AF and thermal demagnetization results, and the major polarity
boundary that divides magnetozones R1 and N2 at 227.16 m is
correlated to the Gauss-Matuyama boundary (G/M: 2588 ka, the
base of the Quaternary) in the GPTS (Gradstein et al., 2012) (Fig. 7a,
b). The two relatively thick magnetozones with positive inclinations
within R1 (n2 and n4) can be correlated to the Jaramillo
(990e1070 ka) and Olduvai (1770e1950 ka) subchrons, respectively
(Fig. 7a, b). Since there are only two subchrons with reversed
polarity in the Gauss chron (Gradstein et al., 2012), magnetozones
r7 and r8 are correlated to the Kanea (3040e3110 ka) and
Mammoth (3220e3330 ka) subchrons, respectively. On the basis of
this correlation, an age of ~3.5 Ma was linearly extrapolated for the
base of the core (Fig. 7a, b). In the Quaternary, the sediments have a
mean sedimentation rate of 8.78 cm/kyr (Fig. 8), which is comparable
to estimated values on 1000-year time scales during the late
Quaternary (Kim et al., 1999). In addition, assuming uniform sedimentation
rates during each period (Fig. 8), some of short-lived
inclination anomalies in the studied core are also tentatively
correlated to the excursions that are well validated (Laj and
Channell, 2007; Roberts, 2008) (Fig. 7a, b).
The reliability of the magnetostratigraphy for core CSDP-1 can
be further assessed by comparison with nearby cores. Regardless of
the variable thickness of Holocene deposits, which range from ~17m
in QC2 (Zhou and Ge, 1990), ~2-m in EY02-2 (Ge et al., 2006),
~4.5-m in NHH01 (Liu et al., 2014a), 0-m in DLC70-3 (Mei et al.,
2016), and ~3-m in CSDP-1, the thickness of sediments deposited
since the mid-Pleistocene (0.78 Ma) in these cores are consistently
60e70 m (Fig. 9). This may not only indicate the reliability in
recording of the M/B boundary but it also suggests a similar
J. Liu et al. / Quaternary Science Reviews 144 (2016) 1e158
sedimentary history among the studied sites during this period. In
addition, in view of the similar position and water depth of cores
CSDP-1 and QC2 (Fig. 1b), correlation of the normal polarity
magnetozone at around 112 m in core CSDP-1 (n2 in Fig. 7b) to the
Jaramillo subchron further indicates that a much older basal age
might also be present in core QC2 (Liu et al., 2014a).
Fig. 6. Orthogonal vector plots and remanence decay curves for (a,c) stepwise AF and (b,d) thermal demagnetization data for representative greigite-bearing samples. The green
dashed line in the remanence decay curve in (a) also denotes the median destructive field (MDF). Acquisition of a GRM is observed at high static AFs (grey bars in a, c), and hightemperature
components due to magnetite often have contrasting polarities to those due to greigite (b,d). (For interpretation of the references to color in this figure caption, the
reader is referred to the web version of this article.)
J. Liu et al. / Quaternary Science Reviews 144 (2016) 1e15 9
With this new chronological framework for the core CSDP-1, the
following issues can be clarified for the first time. First, the continental
shelf of the East China Sea is the northernmost part of the
Chinese continental shelf in which the base of the Quaternary has
been recovered, and the bases of the Quaternary in the Yellow Sea
and Bohai Sea have remained unknown for a long time (He, 2006).
CSDP-1 is also the first core to recover the base of the Quaternary in
the Yellow Sea, which will greatly promote studies of the evolution
of the Yellow Sea on longer time scales.
The second issue concerns the occurrence time and extent of the
first transgression of the Yellow Sea. According to geophysical data,
the Fujian-Lingnan Uplift (FLU), which lies between the southeastern
Korean Peninsula and the Yangtze Estuary (see the
boundary dividing the SYS and East China Sea in Fig. 1b), has
blocked the East China Sea from intruding into the Yellow Sea
(Wageman et al., 1970). The FLU was reported to have subsided in
the early Pleistocene, which allowed sea water ingression into the
Yellow Sea, and several transgressions have subsequently occurred
(Qin et al., 1989). Nevertheless, the details of these transgressions
have not been well investigated due to the lack of chronological
data (Qin et al., 1989).
Our study reveals that the upper 85-m of the studied core
Fig. 7. Magnetostratigraphy of core CSDP-1. (a) Geomagnetic polarity time scale (GPTS) (Gradstein et al., 2012), the validated excursions marked with blue are from Laj and Channell
(2007) and Roberts (2008), and reversals are marked with black. Variations of (b) polarity, (ced) inclinations and maximum angle deviations (MAD), and (e) ratios of saturation
remanence (Mrs) to magnetic susceptibility (c: mass-specific) along the studied core. The black dots in (c) and (d) denote inclinations and MAD values calculated from AF
demagnetization data, while red/blue dots denote thermal demagnetization data after/prior to 400 C. The grey bars mark the greigite-bearing layers in the core, where high Mrs/c
values are taken to indicate SD-like properties due to greigite (Snowball, 1991; Roberts and Turner, 1993). (For interpretation of the references to color in this figure caption, the
reader is referred to the web version of this article.)
J. Liu et al. / Quaternary Science Reviews 144 (2016) 1e1510
Fig. 8. Age-depth relationship for core CSDP-1, with (b) polarity correlating to the (a) GPTS (as shown in Fig. 7a). Only the age points of AMS 14
C dating and major polarity transitions
and subchrons (blue dots) are used to calculate the sedimentation rates (in cm/kyr). (For interpretation of the references to color in this figure caption, the reader is referred to the
web version of this article.)
J. Liu et al. / Quaternary Science Reviews 144 (2016) 1e15 11
contains marine sediments. Three thin (less than 5 m) layers at
depths of around 112 m, 135 m, and 145 m contain greigite, which
were likely to have formed as a result of rapid sea-level change in
the continental shelf sediments (e.g., Oda and Torii, 2004). Therefore,
the occurrence age for the first marine deposit at depth of
145 m is ~1.7 Ma; two short-lived marine transgressions then
occurred at ~1.44 Ma (135 m) and 1.03 Ma (112 m), respectively
(Fig. 10c, d). These early transgressions of the Yellow Sea were
characterized by short durations over small areas (Qin et al., 1989).
Subsequently, seawater retreated until ~0.85 Ma. Then an extended
marine transgression occurred again to create marine environments
in the study area since ~0.7 Ma (Fig. 10c, d).
Studies of nearby core NHH01 (Fig. 1b) indicate that it records
mostly marine environments since ~1.02 Ma (Fig. 10d, e) (Liu et al.,
2014a), which along with core CSDP-1 indicate that the paleoshoreline
probably lay between these two cores between ~1.02
Ma and ~0.85 Ma. The disparate thicknesses of the Jaramillo subchron
in the two cores could be because core NHH01 was closer to
the basin depocenter (Qin et al., 1989) for longer than core CSDP-1
during this period. In addition, several large amplitude sea-level
regressions occurred in glacial periods such as during MIS 20, MIS
18, and MIS 10 (Fig. 10cee), with seawater even retreating from the
central SYS, which resulted in deposition of terrestrial sediments or
erosion.
The third issue involves the effects of local tectonics on the
depositional history of the SYS. Although climate has undoubtedly
played an important role in evolution of the sedimentary environment
of the Yellow Sea (Qin et al., 1989; He, 2006; Liu et al.,
2010), it can not be the only driving factor in sedimentation. The
lowest stand of sea-level is reported to be about À130 m in the
Quaternary (Fig. 10c) (Miller et al., 2005; Rohling et al., 2009). If
only climate is considered, sediments below depths of ~80 m in
core CSDP-1 and ~60 m in core NHH01 should be marine deposits at
least in interglacials since the first marine transgression into the
SYS. This, however, is not the case (Fig. 10c, e). Therefore, the effects
of local tectonic movements such as basement uplift or FLU after its
first subsidence as Jin and Yu (1982) hypothesized, which warrant
further investigation, should be given greater consideration in
future. For the thicknesses of marine sediments documented in
these cores to be present, the continental shelf must have undergone
substantial subsidence coincident with sedimentation.
6. Conclusions
Systematic rock magnetic analyses on the longest sediment core
(CSDP-1) yet recovered from the South Yellow Sea indicate that the
magnetic minerals in the core consist of magnetite, greigite, and
hematite, which occur either as the major magnetic carrier or they
co-exist. The complexity of magnetic mineral assemblages in these
continental shelf sediments, including remanences recorded by
authigenic greigite, has a considerable influence on determination
of a polarity stratigraphy, which is avoided by thermal demagnetization
above 400 C where no influence of greigite remains. The
variable directions of these remanences further support the ‘magnetic
interaction-induced contradictory magnetization’ mechanism.
The M/B boundary is located at a depth of 73.68 m in the
studied core, and the base of the Quaternary in the Yellow Sea has
been recovered for the first time, which corresponds to a depth of
227.16 m in the studied core. The extrapolated basal age of the core
is ~3.50 Ma. The first marine transgression into the Yellow Sea
occurred no later than ~1.7 Ma, with other early transgressions
characterized by short durations over small areas. Terrestrial
Fig. 9. Magnetostratigraphic frames including variations of (a, g, h, n, q) lithology, (b, f, i, m, p) inclinations, and (c, e, j, l, o) polarity along depths of existing long sediment cores
from the SYS. The (d, k) GPTS is from Gradstein et al. (2012). Data for (aec) core QC2 are from Zhou and Ge (1990), (len) EY02-2 from Ge et al. (2006), (hej) NHH01 from Liu et al.
(2014a), and (oeq) DLC70-3 from Mei et al. (2015). The ages “10 ka” marked with blue near the top of the cores were obtained via 14
C dating. The red dots in (f) and (i) denote
inclinations calculated from thermal demagnetization data, and the black dots in (b, f, i, m, p) denote inclinations calculated from AF demagnetization data.
J. Liu et al. / Quaternary Science Reviews 144 (2016) 1e1512
deposition or erosion occurred in several glacial periods such as
during MIS 20, MIS 18, and MIS 10 when seawater retreated from
the central Yellow Sea. Both climate changes and local tectonics
have controlled the evolution of sedimentation in the Yellow Sea
throughout the Quaternary.
Acknowledgements
We are grateful to the reviewers and the Editor for their
thoughtful and constructive comments and suggestions that
greatly improved the manuscript. We thank the crew of the R/V
Kan407 for their assistance with core drilling. Sincere appreciation
is expressed to Qiongyi Luo, Jie Li, Debo Zhao, Jin Zhang, and
Zhongya Chen for their help with sampling. This work was jointly
supported by the China Geological Survey (No. GZH201100202), the
National Natural Science Foundation of China (No. 41430962), the
Project of Taishan Scholar (No. 41330964), and Project of State
Oceanic Administration, China (No. 908-01-BC15).
Appendix A. Supplementary data
Supplementary data related to this article can be found at http://
dx.doi.org/10.1016/j.quascirev.2016.05.025.
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