The continent–ocean transition at the mid-northern margin of the South China Sea

The continent–ocean transition at the mid-northern margin of the South China Sea

    The continent-ocean transition at the mid-northern margin of the South China Sea Jinwei Gao, Shiguo Wu, Kirk McIntosh, Lijun Mi, Boch...

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    The continent-ocean transition at the mid-northern margin of the South China Sea Jinwei Gao, Shiguo Wu, Kirk McIntosh, Lijun Mi, Bochu Yao, Zeman Chen, Liankai Jia PII: DOI: Reference:

S0040-1951(15)00168-7 doi: 10.1016/j.tecto.2015.03.003 TECTO 126566

To appear in:

Tectonophysics

Received date: Revised date: Accepted date:

25 May 2014 20 February 2015 2 March 2015

Please cite this article as: Gao, Jinwei, Wu, Shiguo, McIntosh, Kirk, Mi, Lijun, Yao, Bochu, Chen, Zeman, Jia, Liankai, The continent-ocean transition at the mid-northern margin of the South China Sea, Tectonophysics (2015), doi: 10.1016/j.tecto.2015.03.003

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ACCEPTED MANUSCRIPT The continent-ocean transition at the mid-northern margin of the South China Sea Jinwei Gaoa, f Shiguo Wu b, * Kirk McIntosh c Lijun Mi d Bochu Yao e Zeman Chena, f Liankai Jiaa, f Key Laboratory of Marine Geology and Environment, Institute of Oceanology, Chinese Academy of Sciences, Qingdao,

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266071, China

Sanya Institute of Deep-Sea Science and Engineering, Chinese Academy of Sciences, Sanya, 572000, China

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Institute for Geophysics, University of Texas at Austin (R2200), 10100 Burnet Road, Austin, Texas 78758-0000, USA

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Exploration Department, China National Offshore Oil Corporation Ltd., Beijing 100010, China

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Guangzhou Marine Geological Survey Bureau, Ministry of Land and Resources, Guangzhou 510075, China

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University of Chinese Academy of Sciences, Beijing, 100049, China

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*Corresponding author: Shiguo Wu, Email: [email protected], Tel: +86 898 88380961, Fax: +86 898

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Abstract

The northern margin of the South China Sea (SCS) has particular structural and stratigraphic

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characteristics that are somewhat different from those described in typical passive margin models. The differences are attributable to poly-phase tectonic movements and magmatic activity resulting from the interaction among the Eurasian, Philippine Sea and Indo-Australian plates. Based on several crustal-scale multi-channel seismic reflection profiles and satellite gravity data across the northern SCS margin, this paper analyses the structures, volcanoes and deep crust of the continent-ocean transition zone (COT) at the mid-northern margin of the SCS to study the patterns and model of extension there. The results indicate that the COT is limited landward by basin-bounding faults near Baiyun sag and is bounded by seaward-dipping normal faults near the oceanic basin in our seismic lines. The shallow anatomy of the COT is characterized by rift depression, structural highs with igneous rock and/or a 1

ACCEPTED MANUSCRIPT volcanic zone or a zone of tilted fault blocks at the distal edge. Gravity modeling revealed that a high velocity layer (HVL) with a 0.8-6-km thickness is frequently present in the slope below the lower crust.

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Our study shows that the HVL is only located in the eastern portion of the northern SCS margin based on

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the available geophysical data. We infer from this that the presence of an HVL is not required in the COT at the northern SCS margin. The magmatic intrusions and HVL may be related to partial melting caused by the decompression of a passive, upwelling asthenosphere, which resulted primarily in post-rifting

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underplating and magmatic emplacement or modification of the crust. Based on this study, we propose that

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an intermediate mode of rifting was active in the mid-northern margin of the SCS with characteristics that are closer to those of the magma-poor margins than those of volcanic margins.

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Keywords: South China Sea, Continent-Ocean Transition, multi-channel seismic data, gravity modeling,

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rifting

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ACCEPTED MANUSCRIPT 1. Introduction As one of the largest marginal sea basins of the western Pacific (LaFond, 1966), the South China Sea

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(SCS) is located at the southeast edge of the Eurasian plate, where it is heavily influenced by the Philippine

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Sea plate and the Indo-Australian plate. From the Cretaceous Period until the present, the SCS area has been shaped by several regional processes: 1) the collision of the Indo-Australia plate and the Eurasian plate to the northwest (Morley, 2002; Tapponnier et al., 1986); 2) the slab-pull of the Mesozoic Proto South

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China Sea in the south (Hall, 1996; Morley, 2002; Taylor and Hayes, 1983); and 3) the subduction and

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compression from the Philippine Sea plate in the east (Li, 1994; Zhou et al., 2002). In this active environment, the SCS area has hosted complex structural, volcanic, and fluid activities, acting as an

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exceptional natural laboratory. The SCS provides a particularly attractive opportunity to study two highly

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extended, conjugate rifted margins, separated by a limited ocean basin.

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Passive continental margins can be divided into two end member models with different characteristics regarding the transition from continental to oceanic crust, which we will refer to as the continent-ocean

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transition zone (COT). Magma-poor margins are generally characterized by detachment faults, rotated fault blocks with seaward dipping faults, serpentinized peridotite ridges, and limited syn-rift magmatism (Manatschal and Bernoulli, 1999; Dean et al., 2000; Hopkinson et al., 2004; Pérez-Gussinyé et al., 2006; Lavier and Manatschal, 2006; Reston, 2007; Fernàndez et al., 2010). Alternatively, volcanic rifted margins include a seaward dipping reflector (SDR) sequence, the expression of subaerial lava flow, the presence of sills and dykes in the sediment and a high velocity layer (HVL) in the lower crust resulting from the voluminous magmatism (Eldholm et al., 1987; Mutter, 1993; Dean et al., 2000; Menzies et al., 2002; Ceramicola et al., 2005; Praeg et al., 2005; Geoffroy, 2005; Schnabel et al., 2008; Hirsch et al., 2009; Fernàndez et al., 2010). The characteristics of the COT at the different margins play a critical role in 3

ACCEPTED MANUSCRIPT understanding the extensional processes and early spreading. The northern margin of the SCS has now been described by many investigators. The COT at the northern margin of the SCS is different from those

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in the two end member margins proposed by Tucholke et al. (2007) (Zhu et al., 2012). According to

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previous studies, the COT at the northern margin of the SCS is marked by the volcanoes over the lower slope (Yan et al., 2001) and is defined as a deep-water domain with a water depth of 1500 m to 3000 m or more (Yan et al., 2001; Wang et al., 2006). The COT here may also include several volcanoes and igneous

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rocks in the upper/middle crust and a high velocity layer in the lower crust (Wang et al., 2006). The COT

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has been divided into a rift depression, a volcanic zone and a zone of small, tilted fault blocks based on multi-channel seismic profiles (Zhu et al., 2012). In sum, its COT not only has rotated fault blocks with

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seaward-dipping normal faults (Zhu et al., 2012), but it also has a 25-km-wide zone with apparent volcanic

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edifices at the lower slope seaward of the Pearl River Mouth Basin (PRMB) (Clift et al., 2001). An HVL is

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also widely interpreted beneath the slope (Nissen et al., 1995a; Yan et al., 2001; Wang et al., 2006). Therefore, whether this margin is a magma-poor margin (Yan et al., 2001; Yan and Liu, 2005; Wang et al.,

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2006; Wei et al., 2011; Ding et al., 2012), volcanic rifted margin (Nissen et al., 1995a; Qiu et al., 2003) or an intermediate form (Clift and Lin, 2001; Franke et al., 2011; Zhu et al., 2012) is still debated. Geologic understanding of the SCS has benefitted from commercial hydrocarbon exploration, including the acquisition of thousands of kilometers of seismic reflection lines and the drilling of hundreds of wells (e.g., Chen et al., 1987; Clift and Lin, 2001; Tyrell and Christian, 1992). The crustal structure of the northern margin of the SCS was revealed by Huang et al. (2005) using the first deep reflection seismic profile. Hu et al. (2009) also imaged the crustal structure in the Pearl River Mouth Basin (PRMB) through other multi-channel seismic profiles. In their studies, basin-bounding faults were shown to cut through the upper crust basement and to possibly penetrate into the lower crust. Obvious crustal thinning occurs across 4

ACCEPTED MANUSCRIPT the northern margin of the SCS, and the Moho depth decreases step-wisely from the continental shelf to the oceanic domain (Huang et al., 2005; Hu et al., 2009). Moreover, Huang et al. (2005) found deep structures

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that they interpreted to be a high velocity layer (HVL) and a possibly subducted segment of Mesozoic

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oceanic crust, whereas Hu et al. (2009) found only a step-wise decreasing Moho depth to the south. In addition, several wide-angle refraction seismic lines have been acquired in scientific experiments in the northern margin of the SCS (e.g., Taylor and Hayes, 1980, 1983, Nissen et al., 1995a; Yan et al., 2001; Qiu

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et al., 2001; Wang et al., 2006; Wei et al., 2011; Pichot et al., 2013). An HVL with varied thickness was

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detected from the ESP and OBS data in the eastern portion of the northern margin of the SCS (e.g., Nissen et al., 1995a; Yan et al. 2001; Wang et al., 2006; Wei et al., 2011; Lester et al., 2014), whereas it was not

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found in the OBH and OBS data in the western portion except the early ESP data (e.g., Qiu et al., 2001;

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Ding et al., 2012; Pichot et al., 2013). Hence, it is important to discuss where the HVL pinches out along

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the northern margin of the SCS and whether this may reflect differences in rifting and break-up. This study focuses on the northern SCS COT by interpreting the structures, especially in the deep

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crust, imaged in several crustal-scale multi-channel seismic reflection profiles (MCS) that have been acquired in recent years. All of these profiles cross the continental margin and extend into the oceanic domain, enabling us to investigate the nature of this still controversial boundary. We also use gravity modeling to supplement our seismic interpretations, especially to track the distribution of the HVL. The underlying motivation of this study is threefold: (1) to identify whether a high velocity layer (HVL) is indicated in the multi-channel seismic reflection profiles (MCS), (2) to discuss the possible origin of HVL and the magma source, and (3) to discuss the deformation pattern of the SCS to provide constraints for possible extensional mechanisms. 2. Geologic setting and seismo-stratigraphy 5

ACCEPTED MANUSCRIPT The South China Sea (SCS) is surrounded by the South China Block in the north, the Indochina Peninsula with extensive strike-slip faults in the west, the Palawan-Borneo Islands in the south, and the

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Luzon Arc in the east, and it has an irregular rhombic shape with a SW-pointing apex (Tapponnier et al.,

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1982, 1986; Huchon et al., 2001; Leloup et al., 2001; Morley, 2002; Clift et al., 2008). The SCS region has experienced multiple tectonic events since the Cretaceous (named the Shenhu, Zhuqiong I, Zhuqiong II, Nanhai, Baiyun, and Dongsha events in the northern margin of the SCS) (Taylor & Hayes, 1980, 1983;

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Zhou et al., 1995). The onset of rifting (Shenhu event) in the proto-South China Margin was estimated to

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start at 65±10 Ma in the latest Cretaceous to Early Paleocene as the existing, Mesozoic convergent margin evolved into a zone of extension (Hinz and Schlüter, 1985; Holloway, 1982; Lee and Watkins, 1998; Taylor

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and Hayes, 1983). The rifting (Nanhai event) continued for more than 30 m.y. until seafloor spreading

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occurred in the Northwest and East sub-basins of the SCS. Although still controversial, seafloor spreading

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in the SCS is currently cited by most authors to occur from ~32 Ma to 15.5 Ma (Li and Song, 2012). The SCS basin ended seafloor spreading due to the initiation of eastward subduction of the East sub-basin at

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the Manila Trench (Arfai et al., 2011) and the collision of the Dangerous Grounds (Nansha Islands) and the Reed Bank (Liyue Bank) with Borneo along the NW Borneo Trough (Nansha Trough) in the south (Wang et al., 2006). A limited tectonic event named the Dongsha event subsequently took place along the northern margin of the SCS accompanied with uplift, faults, and magmatic activity during the Middle to Late Miocene (Su et al., 1989; Rao, 1992; Li, 1993; Chen et al., 2003; Cai et al., 2010). Therefore, the complex geological evolution of the SCS resulted in a variety of structures across the basin, including the abundant extensional structures and limited magmatic activity that developed in the northern margin of the SCS during and after the rifting and seafloor spreading. The full evaluation of these features and their deformational history helps to understand the extensional processes of the SCS. 6

ACCEPTED MANUSCRIPT Our study area lies in the Pearl River Mouth Basin (PRMB) and adjacent oceanic basin, located in the mid-northern margin of the SCS (Fig. 1). The PRMB is a Mesozoic–Cenozoic petroliferous basin that can

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be divided into several tectonic zones according to variations in sediment thicknesses (Fig. 1b) (Li and Rao,

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1994; Shi et al., 2005); from north to south, these are the Northern Fault Terrace, the Northern Subsidence Zone (Zhu I and Zhu III Depressions), the Central Rise Zone (Shenhu Rise, Panyu Swell and Dongsha Rise), the Southern Subsidence Zone (Zhu II and Chaoshan Depressions), and the Southern Rise (Shi et al.,

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2005). The seismic profiles presented here mainly run across the Zhu II Depressions (Baiyun sag and

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Liwan sag) and the Southern Rise and extend into the oceanic basin (Fig. 1b). A detailed stratigraphic description has been developed for the Cenozoic section of the PRMB for

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hydrocarbon exploration (Ru and Pigott, 1986; Wu, 1994; Wang and Chen, 1999; Lüdmann and Wong,

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1999) and is presented in Fig. 2. The sedimentary environment in the northern margin of the SCS varied

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through time from terrestrial to marine. Primary regional seismic reflectors (T30 to T70) have been identified and divided into seven seismic sequences characterized by reflection features, seismic facies and

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the presence of gas (Lüdmann and Wong, 1999). 3. Data and methods

The multichannel seismic (MCS) data used in our study were acquired by the South China Sea Institute of Oceanology, Chinese Academy of Sciences (SCSIOCAS) and the China National Offshore Oil Corporation (CNOOC) from 2002 to 2011. The MCS data of line L1 were acquired by SCSIOCAS in 2011 with 144 channels (~1900 m far offset) and were shot using two arrays of Bolt air guns with a total volume of 5080 in3. The channel interval was 12.5 m and the shot interval was 20 s at a speed of 5 knots, resulting in a shot interval of 50 m. The data were sampled at 2 ms and recorded up to 12 s (two-way travel time, hereafter TWT). The MCS data of 7

ACCEPTED MANUSCRIPT lines L2 to L4, some of which have been described by Hu et al. (2009), were acquired by CNOOC from 2002 to 2004 with 576 channels (~200-7400 m offset). In addition to these lines, some petroleum industry

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seismic profiles with record lengths of 5 s to 8 s (TWT) are used to interpret the landward structures where

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these MCS data do not reach. A Moho discontinuity was identified from interpretations of profiles L2 and L4 by Hu et al. (2009). However, in many areas, deep crustal reflections above the interpreted Moho are

and may mark the HVL (Lester et al., 2014).

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also visible. These reflections are similar to deep events interpreted in previous studies farther northeast

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The latest satellite gravity data, compiled in a 1-minute grid, were obtained from the Scripps Institution of Oceanography, University of California San Diego, USA. Combining new radar altimeter

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measurements from the CryoSat-2 and Jason-1 satellites with existing data, Sandwell et al. (2014)

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constructed a global marine gravity model that is two times more accurate than previous models. We used

the deep structures.

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these data to perform 2-D gravity modeling along MCS profiles (L2, L3 and L4) to determine the nature of

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To perform the gravity modeling, a time-depth conversion of the MCS data is necessary to develop a starting depth model. A time-depth relationship was obtained by fitting the time-depth data from well PY33-1-1 with a power function (i.e., D=a*tb+c) (Hu et al., 2009). Hu et al. (2009) noted that this type of function has the advantage that it can prevent the converted depths from being under- or over-estimated by cubic or quadratic polynomial time-depth functions, and it is applied to sedimentary units only. This time-depth relationship was also adopted successfully by Chen (2013) and Xie et al. (2014) in their studies. We used this power function to convert the sediment units to depth. An average velocity of 6.45 km/s proposed by Christensen and Mooney (1995) and 7.3 km/s proposed by Huang et al. (2005) was used to estimate the thickness of the crust and the interpreted HVL, respectively. As a test, we compared the Moho 8

ACCEPTED MANUSCRIPT depths of our depth model with those of the depth model of Hu et al. (2009) and the OBS model of Yan et al. (2001) in our study area (Fig. 3). The results show that the Moho depth of profile L2 in our study is a

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slightly deeper than that of Hu et al. (2009) generally, and profile L4 has a similar Moho position

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compared to the model of Hu et al. (2009). Compared with the OBS model of Yan et al. (2001), both depth models of profile L4 may under-estimate the Moho depth, whereas the Moho depth of profile L2 is possibly over-estimated in the shelf and slope. Although there are some differences, these models are all

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quite similar.

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Using the depth-converted seismic sections, we developed a simple four-layer crustal model to approximate the crustal structure, that is, sea water, sediments (including several sedimentary layers), crust

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(including continental and oceanic crust) and mantle. The initial densities shown in Table 1 were adopted

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from previous investigations performed in the SCS (Kido et al., 2001; Trung et al., 2004; Tsai et al., 2004;

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Braitenberg et al., 2006; Hao et al., 2008, 2011; Franke et al., 2011). Considering the offline effects not included in 2-D modeling, the deviations between the calculated and observed gravity are due to the

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uncertainties of the modeling. 4. Analysis and results

4.1 Terrigenous clastic sediments extending to the oceanic basin An igneous body at a distance of 255-270 km along seismic profile L1 that prevented the terrigenous clastic sediments from extending to oceanic basin has been interpreted by Li et al. (2008) (Fig. 4). However, new results from the IODP leg U1435 of Expedition 349 show that there is an obvious unconformity (~33 Ma) in this „igneous body‟ that separated the marine sediments from the poorly sorted littoral sandstone and black mudstone (Expedition 349 Scientists, 2014). Therefore, this „igneous body‟ is now known to be a structural high rather than a volcano (Fig. 4), indicating that similar bathymetric highs 9

ACCEPTED MANUSCRIPT below the lower slope of profiles L2 to L4 may be Mesozoic to Paleogene sedimentary rock rather than igneous rock (Figs. 5 to 8).

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Another important observation is that Reflector T60 marks the interface between the Oligocene and

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Miocene with a typically high-amplitude seismic reflection in PRMB (Figs. 5 to 8). This event corresponds to an age of approximately 23.8 Ma and is associated with a spreading ridge jump to the south (Shipboard Scientific Party, 2000; Franke et al., 2013). As an important unconformity, Reflector T60 also represents

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the boundary between the lower terrigenous clastic facies and overlying marine facies.

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Reflector T60 in the northern SCS has previously been interpreted to terminate on the lower slope (e.g., Hu et al., 2009; Chen, 2013). However, our seismic interpretations show that Reflector T60 crosses

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the lower slope and pinches out near the ocean-ward zone of tilted fault blocks at distances of

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approximately 280 km, 260 km and 220 km along profiles L2, L3 and L4, respectively, indicating that the

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deposits of terrigenous clastic sediments likely crossed the basement high ocean-ward and extended tens of kilometers farther toward the oceanic basin rather than pinching out in the lower slope (Figs. 5 to 8).

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4.2 Deep structures

In the deeper portions of our seismic profiles, we have interpreted two sets of seismic reflectors (named R1 and R2) (Figs. 5, 6, 7 and 9). Reflector R1 is interpreted as the seismic reflection marking the Moho, or base of the crust. R1 is characterized by relatively continuous, high-amplitude, low-frequency seismic reflections within the range of 9-11 s (TWT) below the lower slope and dipping landward (deepening to the north) in profiles L2, L3 and L4 (Figs. 5, 6, 7 and 9). R1 reflections are also visible below the oceanic basin at 8-8.5 s (TWT), which is shallower than that below the lower slope. We observe an abrupt change of R1 reflections at distances of 280 km, 260 km and 220 km along profiles L2, L3 and L4, respectively, which we interpret to be the location of the continent-ocean boundary (COB; south of this 10

ACCEPTED MANUSCRIPT is true oceanic lithosphere) between lower slope and oceanic basin (Figs. 5 to 7). There is another deep reflection visible along profiles L2 and L3 in a position above reflector R1.

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This reflector, R2, is characterized by medium- to high-amplitude seismic reflections but relatively less

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continuous events than that of reflector R1 (Figs. 5, 6 and 9). It occurs between 8 s and 10 s (TWT) below the lower slope and dips generally landward, sub-parallel to R1 events (Figs. 5, 6 and 9). The thickness of the layer between reflectors R1 and R2 varies from 0.25 s to 1.62 s (TWT), which corresponds to

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approximately 0.8-6 km in thickness. A high velocity layer (HVL, 7-7.5 km/s) with a thickness of 4 km

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(average) and estimates of 3-12 km and 0-5 km at the base of the crust was reported below the lower slope along Profiles OBS1993, OBS2006-1, and OBS2001, respectively (Yan et al., 2001; Wei et al., 2011; Wang

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et al., 2006). A thick HVL was also interpreted below the slope along Profile ESP-E (Nissen et al., 1995a).

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Similar to the events described by Lester et al. (2014), we suggest that reflector R2 is likely to be a seismic

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reflection interface at the top of an HVL. This is also consistent with the interpretation of Sun et al. (2009) in their seismic profile. We note that no clear R2 seismic reflection interface is observed in the similar

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position below the lower slope along profile L4. This may imply that the HVL does not occur in the area to the west (Figs. 7). Below, we use gravity modeling to test the suggested relationship between R2 reflections and an HVL.

Another medium- to high-amplitude, relatively continuous seismic reflection (designated R3) is identified below the lower slope along profile L3 (Figs. 6 and 9). This event is at 6 s to 8 s (TWT), is between reflector R2 and the top of the basement (Tg), extends from the structural high to the basement high, and disappears at the transition to oceanic crust. This interface nearly merges with the basement reflection (Tg) at a distance of 168-172 km along profile L3 (Fig. 6). We suggest that this rough seismic reflection interface may represent a boundary (perhaps a detachment fault) between the upper and lower 11

ACCEPTED MANUSCRIPT crust. 4.3 Gravity modeling

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The gravity models created along seismic profiles L2, L3 and L4 are presented in Figs. 10, 11 and 12,

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respectively. The landward boundaries for the Moho and the top of HVL are inferred from the gravity data only because corresponding reflections could not be interpreted in this part of the seismic profiles. Several necessary steps are used to perform the gravity modeling as suggested by McIntosh et al. (2014).

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From the initial model of profile L2, although the long wavelength variations of the calculated gravity

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are well correlated with the Moho depth, there exists significant misfit with the observed gravity (Step I, Figs. 10a and d). Thus, we assigned a density of 2.97 g/cm3, corresponding to a seismic velocity of 7-7.5

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km/s, and the gabbroic lithology to the HVL identified in seismic profile L2 (Step II, Figs. 10b and d). This

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step significantly improved the fit, but there was still an offset between the calculated and the observed

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gravity. We noticed that an uppermost mantle layer with a velocity of 8.0-8.3 km/s and a thickness of 8-12 km was revealed by Yan et al. (2001) in their OBS model and that a similar upper mantle layer with a

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thickness of 10 km was presumed for gravity modeling in the northeast SCS (McIntosh et al., 2014). In our final model, we also included an uppermost mantle layer with a thickness of 10 km or so and assigned a density of 3.0-3.15 g/cm3 to this layer; this significantly improved the calculated gravity fit to the observed data (Step III, Figs. 10c and d). Similarly, the gravity modeling derived from seismic profile L3 was performed as described for profile L2 (Fig. 11). Because the R2 reflection was not recognized from seismic profile L4, we performed the modeling steps as described for profiles L2 and L3, except we assigned a density of 2.97 g/cm3 to a lower crustal body during the gravity modeling (Steps I and II, Figs. 12a, b and d). Although the calculated gravity fits the observed gravity generally, there was still an offset in Baiyun sag, where the calculated 12

ACCEPTED MANUSCRIPT gravity was obviously lower than observed data (Step II, Figs. 12b and d). Therefore, we added a local layer with a thickness of 2-4 km and a density of 2.97 g/cm3 below the depression in the base of the crust.

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With this modification, the calculated gravity fit the observed data well (Step III, Figs. 12c and d),

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suggesting that there may be a small HVL below the depression (Baiyun sag) in this profile. In addition, a general observation is that the observed and calculated gravity data decrease southeastward gradually from high values in the upper slope to low values in the lower slope. They then

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increase gradually to high values again in the oceanic basin. Following the modeling of this study, this

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anomaly pattern appears to be related to significant changes from continental crust to thinned and subsided transitional crust and then to normal oceanic crust.

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5. Discussion

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5.1 COT at the mid-northern margin of the SCS

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According to interpretation of early magnetic and gravity anomalies, the COT here was delineated through abrupt change in crustal thickness and was marked by a negative free air gravity anomaly in the

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northern SCS (Taylor and Hayes, 1983). The oceanic crust in the northern SCS has generally been interpreted to be present only south of approximately 20°N based on magnetic anomalies (Taylor and Hayes, 1980; Briais et al., 1993). North of this, a velocity profile of OBS1995 across the Hengchun Peninsula off SW Taiwan shows that 11-16-km-thick transitional crust subducts eastward beneath the Manila Trench (McIntosh et al., 2005, 2013). Recent seismic reflection and wide-angle data indicate that the transitional crust at the northeastern margin of the SCS is highly extended with typical thicknesses of 7-15 km and is even hyper-extended (here, <4 km thick) in places (Van Avendonk et al., 2009; Pichot et al., 2013; Lester et al., 2014). Using two-ship expanding spread profiles (ESP), Nissen et al. (1995a) found that the crust shows a 13

ACCEPTED MANUSCRIPT general trend of seaward thinning toward the COB and identified an HVL in the lower crust along the northern SCS margin. Two velocity profiles based on OBS data recorded in the northern SCS also

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indicated the nature of the COT below the lower slope (Yan et al., 2001, OBS1993; Wang et al., 2006,

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OBS2001). In these profiles, the COT was characterized by volcanoes and igneous rocks in the upper/middle crust, an HVL with a P-wave velocity of 7-7.5 km/s at the base of the crust, crustal thinning seaward from the continent to the oceanic basin (22 km to 9 km, OBS1993; 15 km to 11 km, OBS2001)

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by Wang et al. (2006) using these data (Fig. 1a).

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and the Moho depth (12-22 km of OBS 1993, 11-15 km of OBS 2001). An extent of COT was delineated

Farther west, an ocean-bottom hydrophone (OBH) seismic profile deployed in 1996 across the Xisha

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Trough shows crustal thinning from 25-km thick to 13-km thick without evidence of an HVL present (Qiu

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et al., 2001). Moreover, two velocity profiles across the Pearl River Mouth Basin (PRMB) based on OBS

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data recorded in 2006 (Location in Fig. 1: OBS2006-1 nearby Zhongsha Islands (ZI), OBS2006-3 nearby Dongsha Islands (DI)) were recently analyzed to determine the crustal structure (Ding et al., 2012; Wei et

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al., 2011). The OBS2006-3 profile indicates an HVL in the lower crust, whereas the OBS2006-1 profile does not have evidence of such a layer at the base of the crust. Using the multichannel seismic profiles acquired in 2010, Zhu et al. (2012) divided the shallow structures of the COT at the northern margin of the SCS into three morphological units: the rift-depression, the volcanic zone, and a seaward zone of tilted fault blocks, and also delineated as a domain of COT (Fig. 1a). In light of these recent advances from MCS and wide-angle seismic data, we suggest that the following characteristics should be used to identify the configuration and nature of the COT at the northern margin of the SCS: (1) Crustal thickness and the Moho depth change from the continental to oceanic basin. 14

ACCEPTED MANUSCRIPT (2) Volcanoes and igneous intrusive rocks may be present in the lower slope resulting from magmatic upwelling.

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(3) An HVL is likely to be present in the lower crust.

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(4) A rift-depression and tilted fault blocks associated with stretched continental crust are expected. (5) Distinctive gravity and magnetic anomaly patterns associated with presumed oceanic crust, transitional crust, and continental crust and an HVL, if present.

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(6) The lateral extent of the break-up unconformity is limited ocean-ward to the COT.

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Gravity modeling of the crustal structures along seismic profiles L2, L3 and L4 shows a southeastward-tapered thinning of the crust from a thickness of approximately 25 km below the shelf to

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5-6 km near the oceanic basin (Figs. 10 to 12), which means that the stretching factor β is within the range

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of 1.2-6 (assuming the pre-rift crust is 30 km). Crustal thinning southeastward along these profiles also

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results in the southeastward step-wise shallowing of Moho depth from approximately 25 to 26 km below the shelf to 10 to 11 km below the oceanic basin. This observation is similar to previous investigations in

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the northern margin of the SCS (Nissen et al., 1995a; Yan et al., 2001; Huang et al., 2005; Wang et al., 2006; Hu et al., 2009; Wei et al., 2011; Ding et al., 2012; Lester et al., 2014). According to the seismic-constrained gravity modeling (Figs. 10 to 12), the Baiyun sag is a rift depression with rotated fault blocks and is classified as highly extended continental crust with a sharp thickness change from 20 to 24 km below the shelf to 5 to 6 km below the depocenter of the sag. These thickness changes indicate a stretching factor β in the landward edge of the sag of 1.3-1.5 and up to 5-6 in the center of this rift depression, similar to the results of the studies of Hu et al. (2009) and Huang et al. (2005). All these results show that Baiyun sag has experienced focused stretching during rifting. Moreover, some igneous sills are imaged in the seismic profiles in the Baiyun sag and near the structural high as 15

ACCEPTED MANUSCRIPT bounded by large-scale normal faults (Fig. 13). A typical saucer-shaped sill with a fold can be observed at the seaward side of the structural high along seismic profile L4 (Fig. 13b). Thick Early Miocene igneous

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rocks, mainly including tuff lava, breccia lava and lava, have been found in drill cores of BY7-1-1, located

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in the Baiyun sag (Pang, 1988; Yan et al., 2001). The presence of these rocks indicates that Baiyun sag was also affected by magmatic intrusions. In addition, a previous seismic refraction profile shows that an HVL below the lower crust went across this sag (Yan et al. 2001). Both observed and calculated gravity data

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show an obvious change from high to low gravity between the shelf and this sag (Figs. 9 to 11). Therefore,

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based on the criteria listed above, the landward side of the COT is located in the Baiyun sag. Syn-extension sediments, which can be considered a sign of the COT, have been studied in depth in

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the Gulf of Aden by several authors (e.g., d‟ Acremont et al., 2005, 2006; Autin et al., 2010; Leroy et al.,

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2010; Watremez et al., 2011). They found that the occurrence of the syn-extension sediments is

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contemporaneous with the deformation of the basement in the COT and does not exist on adjacent oceanic crust, i.e., deposition precedes break-up. They suggested that these sediments were most likely related to

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the upper syn-rift sequence and were covered by post-rift deposits as were the upper syn-rift series. Similar characteristics also can be observed in our seismic profiles (Fig. 8). The lower terrigenous clastic sediments were deposited on the deformed basement and pinched out near the ocean-ward zone of tilted fault blocks. Therefore, we suggest that the lower terrigenous clastic sediments bounded by reflectors Tg and T60 between the lower slope and oceanic basin be considered the syn-extension sediments and that the COB extends farther than the ocean-ward side of the lower slope. The limit of this unit does correspond to the abrupt change of the Moho depth, the disappearance of the HVL, the edge of the thinnest crust (oceanic crust) and an obvious seaward change from low to high gravity (Figs. 10 to 12). Based on these observations and the criteria above, we prefer to adopt the extent of the COT as suggested by Wang et al. 16

ACCEPTED MANUSCRIPT (2006) rather than that of Zhu et al. (2012) in the mid-northern margin of the SCS. According to our analysis of the seismic profiles, the shallow anatomy of the COT with an extent of

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245 km along seismic profile L1 is characterized by rift depression and structural highs with igneous rock

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bounded by normal fault (Fig. 4). The shallow characteristics of the COT with an extent of 240-265 km along seismic profiles L2 and L3 are similar to seismic profile L1, including rift depressions, structural highs and tilted faults blocks (Figs. 5 and 6). However, the shallow characteristics of the COT with an

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extent of 220 km along seismic profile L4 are slightly different from those of other seismic profiles. It is

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marked by rift depressions, a structural high affected by lava flow, a volcanic zone, and a zone of tilted faults blocks, bounded by normal faults at both sides of the COT (Fig. 7). Commonly, a seaward-dipping

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normal fault bounded the tilted fault blocks of the COT to the oceanic basin. However, the tilted fault

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blocks in our seismic profiles are not well organized and are less obvious than those of Zhu et al. (2012).

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5.2 Extent and origin of High Velocity Layer To examine the distribution of a deep crustal high-velocity layer along the northern margin of the SCS,

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we have plotted simplified profiles from published works that have identified an HVL (Figs. 14 and 15). There are nine primary ESP, OBH and OBS profiles across the northern margin of the SCS (Fig. 1a), with some additional profiles in the Manila Trench area not described here. Nissen et al. (1995a) first identified the HVL as 7-7.5 km/s below the lower crust of the northern SCS using ESP data. Details of this interpretation, especially the very large HVL thickness beneath the upper slope and shelf, have since been questioned due to inadequate deep data quality (Figs. 14d and i) (Yan et al., 2001). However, the HVL was imaged with OBS data and interpreted to be 4 km thick below the shelf and slope along Profile OBS1993 (Yan et al., 2001; Fig. 14f). Moreover, a thicker HVL, 3-12 km thick, below the shelf and slope in the lower crust was reported along Profile OBS2006-3 in the mid-northern margin, suggesting a thickening 17

ACCEPTED MANUSCRIPT tendency eastward (Wei et al., 2011; Fig. 14e). However, a thinner HVL, 0-5 km thick, was interpreted to be present below the lower slope along Profile OBS2001 in the northeastern margin, implying that the

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HVL may pinch out farther eastward (Wang et al., 2006; Fig. 14c). In contrast, Profiles OBS1995 in the

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northeasternmost margin and Profiles OBS2006-1 and OBH1996 in the northwestern margin did not display the presence of an HVL in the lower crust (McIntosh et al., 2005; Qiu et al., 2001; Ding et al., 2012; Figu.14a, j and h). Recent results from the TAIGER project showed clear evidence for a ~3-5-km-thick

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HVL over a large region just west of the Manila trench (Lester et al., 2014; Fig. 14g) and a much thinner

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HVL close to the Manila trench, according to the MCS data interpreted by McIntosh et al. (2014). These results mean that the HVL had been underthrusted at the Manila Trench with transitional crust. Overall,

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the northern SCS margin.

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these profiles indicate that the HVL is widely distributed, although preferentially in the eastern portion of

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In addition to the transects noted above, an onshore-offshore project to jointly record deep seismic refractions and reflections using land seismic stations and OBSs was performed by the South China Sea

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Institute of Oceanology, Chinese Academy Sciences (SCSIOCAS) offshore of Hong Kong in 2004 (OOS2004, Xia et al., 2010; Figs. 14b and 15). This study did not detect the 7-7.5-km/s HVL at the base of the crust, which means that the HVL along Profiles OBS1993 and OBS2006-3 to the southeast does not extend to the nearshore area or onshore. Our seismic sections and gravity modeling show that the HVL continues to extend westward from Profile OBS1993 and terminates near seismic profile L4 under the upper slope. Using all these data, we determined a rough extent of the HVL, which is marked by the white outline in Fig. 15. The HVL, suggested by Franke (2013) to be limited most likely in the northeastern margin of the SCS, is mainly located in the eastern portion of the northern SCS, including the shelf and the slope, and tapering 18

ACCEPTED MANUSCRIPT southeastward to the oceanic basin. Wei et al. (2011) indicate that the HVL is up to 12 km thick beneath the Dongsha Rise along Profile OBS2006-3 and is the thickest occurrence based on OBS profiling. From

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the Dongsha Islands, the HVL gradually thins to both sides below the lower slope and finally terminates at

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the Zhu II depression (Baiyun sag) of the PRMB and the Manila Trench. We also note that the HVL is distributed essentially within the COT but is clearly not present everywhere in the COT. Thus, our investigation shows that the HVL should not be considered as a necessary sign of the COT at the northern

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margin of the SCS.

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In addition to BY7-1-1, some exploration wells and scientific drill holes have revealed that the ages of igneous rock varies from rifting to post-rifting and post-seafloor spreading (PY16-1-1, 41.2±2.0 Ma;

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LH11-1-2, 27.2±0.6 Ma; ODP1148, <1 Ma; see locations in Fig. 1) (Yan and Liu, 2005). Furthermore, the

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igneous sills developed between Reflectors T60 and T40 at the seaward side of the structural high along

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seismic profile L3 (Fig. 13a), whereas they developed on both sides of the structural high along seismic profile L4 between Reflectors T70 and T60 (Fig. 13b). These results imply that the some magmatic activity

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could have occurred in the syn-rift period as well as during and after the seafloor spreading episode of the SCS, consistent with the observed results from drilled wells. A narrow zone (25 km wide) of volcanic rocks was identified under the seaward part of the lower slope seaward in the PRMB by Clift et al. (2001), while Zhu et al. (2012) imaged a 25-45-km-wide volcanic zone within the COT in the PRMB and a much wider zone of volcanic rocks was identified by Wang et al. (2006) in the northeastern SCS. However, Yan et al. (2006) proposed that this volcanic zone is a “patchy band” instead of a continuous band based on a summary of observations of magmatism along the northern SCS. Similarly, a small volcano and a 60-km-wide volcanic zone within the COT in the PRMB are imaged along seismic profiles L1 and L4 (Figs. 4 and 7). The interpretation of volcanism along profiles OBS1993 and OBS2001 indicates that they 19

ACCEPTED MANUSCRIPT cause variation of the sediment thickness beneath the lower slope and the oceanic basin, suggesting that the intrusive/extrusive magmatism and the HVL across the COT occurred primarily after seafloor spreading

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(Yan et al., 2001; Wang et al., 2006; Clift et al., 2001; Zhu et al., 2012). This means that magmatism was

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more active during the Middle or Late Miocene to recent time along the northern margin of the SCS. In addition, because almost no volcanoes or late-stage igneous rocks were identified along OBH1996 and OBS2006-1 in the northwestern margin of the SCS (Qiu et al., 2001; Ding et al., 2012), it appears that the

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magmatism below the lower slope in the northeastern SCS has been more active than in the northwestern

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SCS.

Although most authors have linked the HVL to relatively late-stage magmatism as discussed above,

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there are three candidates that have been suggested for the origin of the HVL in the SCS: 1) residual mafic

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rocks inherited from pre-rifted crustal structure (Nissen et al., 1995a), 2) post-rifting underplating related

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to the volcanism during or after the cessation of seafloor spreading of the SCS (Yan et al., 2001; Wang et al., 2006) and 3) a serpentinized layer at the upper mantle (Franke et al., 2011; Pichot et al., 2013). The

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first candidate may be ruled out because there is no HVL with a velocity of 7-7.5 km/s detected at the base of the crust along Profile OOS2004 offshore of Hong Kong (Xia et al., 2010) (Fig. 14b) and because of the limited extent of the HVL (Fig. 15). Furthermore, serpentinization is generally believed to be due to the interaction of the mantle peridotites with seawater through faults that link the upper mantle to the seafloor (Pichot et al., 2013). Thin HVLs (2-3 km) are frequently interpreted as serpentinized upper mantle and are often found at the necking zones and COTs (Pichot et al., 2013). However, the thickness of the HVL in the northern margin of the SCS was reported to be an average of 4 km and 0-5 km in the lower crust below the lower slope along Profiles OBS1993 and OBS2001 (Yan et al., 2001; Wang et al., 2006) and even 3-12 km near the Dongsha Rise along Profile OBS2006-3 (Wei et al., 2011). Our final model shows that an HVL at 20

ACCEPTED MANUSCRIPT the base of the crust is distributed across the slope with a thickness of 0.8-6 km. Because the extent of the HVL is greater than that of the COT from continent to ocean and because the seawater should have to

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percolate through approximately 10-16 km of transitional crust to reach the upper mantle, it seems difficult

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to explain the origin of the HVL as resulting from serpentinization, which effectively occurs at a maximum depth of 3-4 km below the sea-floor (Cannat et al., 1992, 1997, 2006).

The thicknesses of underplated material in general is typically in excess of 5-6 km (Watremez et al.,

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2011; Pichot et al., 2013), which corresponds to the thicknesses of the HVL reported along the northern

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margin of the SCS. A previous study argued that an HVL due to underplating below the lower crust, which has a lower density than mantle material, would result in buoyancy-driven uplift (Watts, 2001). Our study

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shows that the thickest interpreted HVL (12 km) below the lower crust is located near the Dongsha Islands,

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which is also mainly near the center of our mapped area of the HVL (Fig. 15). We also observe that the

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HVL thins to both sides gradually from the Dongsha Islands and finally terminates at the Baiyun sag of the PRMB and the Taixinan Basin (TXNB) or the Manila Trench, corresponding to structural change from the

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uplift to depressions. Furthermore, our gravity modeling required that an uppermost mantle layer with a varied density of 3.0 to 3.15 g/cm3 and a thickness of 10 km be added to match the calculated gravity and the observed gravity. Although we cannot rule out serpentinization below the oceanic crust where the seawater may percolate through the thin faulted oceanic crust with a thickness of 5-6 km to the underlying upper mantle, percolation below 10-16-km-thick crust may be less likely without clear fluid pathways to the mantle for hydration (McIntosh et al., 2014). Furthermore, some necking models in the northern SCS suggested that if the HVL belonged to the initial crust, the shape of the basement could be very different from the observed basement, especially in the case of a strong lithosphere (Shi et al., 2005). However, if the HVL was not an original part of the crust, it would be similar to the observed basement even in the case 21

ACCEPTED MANUSCRIPT of a strong lithosphere (Shi et al., 2005). Therefore, the underplating appears to be a reasonable explanation for the HVL. However, we noticed that previous heat flow versus subsidence modeling

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showed that an underplated HVL may be inconsistent with the observed modest subsidence (Nissen et al.,

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1995b). Taking into account that the magmatic intrusions/extrusions occurred in the northern SCS, especially in Dongsha Rise, a magmatic emplacement or modification of the crust may also affect the subsidence in addition to the underplating.

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The Cenozoic igneous rocks in the northern SCS may result from the uplift of the thinned crust due to

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the lithospheric extension and seafloor spreading or upwelling of the asthenosphere when the proto-SCS margin began to break-up during the Middle Oligocene (Holloway, 1982). Alternatively, they may be

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related to the lower lithospheric delamination (Zou et al., 1995), or they may represent a crustal weak zone

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emplaced by the deep mantle materials (Zhu et al., 2012). The igneous rocks have also been inferred to be

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related to upper mantle upwelling into the rifted crust and flowing southeastward through fractures and faults in the upper/middle crust after the end of the seafloor spreading (Wang et al., 2006). Similarly, the

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underplating materials have also been proposed to originate from partial melting caused by the decompression of passive, upwelling asthenosphere, an idea that has been applied to the rapidly rifted volcanic margins (Yan et al., 2001). However, despite the observed magmatism, the northern margin of the SCS is different from known volcanic margins because of its slow seafloor spreading (~3 cm/yr, half the seafloor spreading rate) and lack of seaward dipping reflector sequences (Yan et al., 2001). A high heat flow zone with an average surface heat flow of 90 mW/m2 was found in the COT at the northern margin of the SCS (Nissen et al., 1995b; Shi et al., 2003), which is higher than the heat flow results of ODP Leg 184 at Site 1148 (69 mW/m2) and slightly lower than the preliminary heat flow value of IODP 349 at Hole U1432 (94 mW/m2) (Fig. 1b). Moreover, combining surface heat flow values with a 22

ACCEPTED MANUSCRIPT variation of 56-101 mW/m2 acquired by Li et al. (2010) in the COT with the preliminary results of the Expedition 349 Preliminary Report (Expedition 349 Scientists, 2014), we see an increasing trend of heat

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flow values from continent to ocean. In general, the heat flow values gradually increase from the

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continental shelf to the oceanic basin with an anomalous high heat flow zone in the COT. This temperature difference from the continental shelf and oceanic basin may be related to small-scale convection occurring during and after rifting, which is similar to the study of Lucazeau et al. (2008, 2009), who studied the

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rifting processes in the ultraslow spreading (1.3 cm/yr) of the eastern Gulf of Aden. Therefore, we suggest

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the following scenario to explain our observations. At first, the magmatic volume resulting from partial melting due to the decompression of passive, upwelling asthenosphere was not enough to intrude the

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crustal weak zone voluminously following the lithospheric extension during the early rifting and seafloor

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spreading of the SCS. Throughout this period, only a limited amount of lava intruded into the syn-rift and

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post-rift sediments through the deep crustal faults developed in the rift zone. After the end of the seafloor spreading of the SCS, the magmatic flow no longer followed the extension. The accumulation of magma

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below the continental crust in the shelf and slope (including COT) may have produced the expected thermal and density anomaly through the advection of hotter and lighter materials (Lucazeau et al., 2008), which led to some of the hotter and lighter materials intruding into the hyper-extended crust (e.g., near Baiyun sag and Liwan sag) through deep crustal faults; some of the magma even extruded at the seafloor. Magma that failed to penetrate the crust pooled at its base to form the HVL or remained in the uppermost mantle, where it caused the modest density reduction indicated by the gravity models. 5.3 Rifting Mode in the mid-northern margin of the SCS Two end-members of passive margins have been characterized on the Atlantic conjugate passive margins (e.g., Iberia-Newfoundland margins, Norway and SE Greenland margins) through the 23

ACCEPTED MANUSCRIPT identification of their typical features, which we mentioned above (Figs. 16a and b) (e.g., Tucholke et al., 2007; Eldholm et al., 1987, 1989; Larsen et al., 1994, 1999). However, this dichotomy cannot reflect the

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global range of rifted margins adequately. This range more likely covers a broad spectrum, from

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magma-poor to volcanic. One margin with intermediate features is located in the northern Gulf of California. Here, moderate to high heat flow (values of ~100-120 mW/m2) and lateral pressure gradients within the crust produced by strong thickness variations in the crust have driven the flow of the weak

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lower crust (González-Fernández et al., 2005). And this margin is characterized by small amounts of

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magmatism with sills, the combination of which forms non-oceanic “new” crust (Sawyer et al., 2007). The West African north-Angolan margin in the central South Atlantic may be another intermediate margin

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between magma-poor and volcanic. This margin is typically characterized by limited syn-rift magmatism,

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evaporates deposited in the shallow water, and an HVL consistent with magmatic underplating at the base

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of the crust (Huismans and Beaumont, 2011). Generally, these two margins are different from volcanic passive margins, whereas they have some common features with the magma-poor margins; this is also the

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case with the northern margin of the SCS. Rift depressions are expected to reach the maximum tectonic subsidence before the beginning of seafloor spreading according to models of Atlantic-type passive margins (e.g., Steckler and Watts, 1978). However, by using the back-stripping method of Sclater and Christie (1980), Dong et al. (2008) calculated the tectonic subsidence rate of the PRMB based on depth-converted seismic profiles. They found that the tectonic subsidence rate of the rift depressions in the PRMB continued to increase after the onset of seafloor spreading in the SCS during 32-23.8 Ma compared with the syn-rift tectonic subsidence rate. Subsequently, the whole PRMB began to undergo a thermal subsidence after 23.8 Ma (Dong et al., 2008). Franke et al. (2013) considered the shallow-water carbonates as an indication of a delay in syn-rift 24

ACCEPTED MANUSCRIPT subsidence over much of the SCS rift in the Early Miocene. Unfortunately, we do not identify direct evidence of shallow-water carbonates in our seismic profiles, but previous studies show that carbonate

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platforms between Reflectors T40 and T50 (Fig. 2) are well developed in the Dongsha Rise area and

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adjacent depressions and are among the important factors that resulted in focused fluid flow (gas chimneys or hydrothermal fluid) in this area (Sattler et al., 2004; Wu et al., 2009; Sun et al., 2012a; Sun et al., 2012b; Chen et al., 2013; Wu et al., 2014). Therefore, based on the subsidence history, we conclude that the

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mid-northern SCS continued to rift, perhaps episodically, after approximately 32 Ma until approximately

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23.8 Ma.

Furthermore, limited magmatic intrusions occurred in the northern SCS following rifting and seafloor

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spreading of the SCS, but, similar to previous investigations (Yan et al., 2001; Wang et al., 2006; Wei et al.,

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2011), we found that more-intensive volcanic activities altered the seafloor topography at the lower slope

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and COT after the cessation of the seafloor spreading of the SCS. Finally, in combination with the discussion about the high velocity layer, we propose an intermediate

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mode to describe the mid-northern margin of the SCS (Fig. 16c). As described below, this type of margin has a different combination of characteristics than the two end member passive margins. These characteristics include the following: (1) wide regions of COT featuring stepwise crustal thinning seaward; (2) un-deformed, thin syn-extension sediments present between the lower slope and oceanic basin; (3) capping of the post-rift sediments with carbonates, which were deposited in shallow water conditions in the slope; (4) no clear evidence of exposed mantle lithosphere but some limited syn-rift magmatism characterized by sills and tuff and significant post-seafloor spreading volcanism at the seaward part of the lower slope; (5) a local extent of the HVL below lower crust consistent with magmatic underplating during “post-rift” events. 25

ACCEPTED MANUSCRIPT Despite advances in understanding the variations in rifting processes globally and within the SCS area, fundamental information is still missing that could help us fully grasp the SCS extension mechanisms and

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history of rifting. Directly associated with this, it is also critical to study the formation processes and

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timing of the magmatism and HVL related to crustal thinning and mantle/asthenosphere upwelling across the lower slope and COT. For example, very little hard information is available on the age, origin and geochemical character of this magmatism. Furthermore, because there may be a continuum between the

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the SCS will need to utilize scientific drilling.

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two passive margin endmembers, future studies focusing on understanding the fundamental processes of

6. Conclusions

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1) We observed that syn-extension sediments crossed the basement high ocean-ward and extended

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tens of kilometers farther toward the oceanic basin, rather than terminating in the lower slope. We prefer to

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adopt the extent of COT at the northern SCS COT divided by Wang et al. (2006), in which the COT is limited by basin-bounding faults in the Baiyun sag, corresponding to crustal thinning, the change of Moho

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depth and an obvious seaward change from high to low gravity. Near the oceanic basin, it is bounded by seaward-dipping normal faults, corresponding to the termination of syn-extension sediments, the abrupt shallowing of the Moho depth, the edge of the thinnest crust (oceanic crust), an obvious seaward change from low to high gravity and/or disappearance of the HVL. Therefore, the shallow anatomy of the COT, with a width of 240-265 km along seismic profiles L1, L2 and L3, is characterized by rift depression, structural highs with igneous rock and/or a zone of tilted faults blocks, whereas the shallow characteristics of the COT, with an extent of 220 km along seismic profile L4, is marked by rift depressions, a structural high affected by lava flow, a volcanic zone and a zone of tilted faults blocks instead. 2) Gravity models show that a high density layer, directly associated with a high velocity layer (HVL), 26

ACCEPTED MANUSCRIPT interpreted based on seismic reflection and refraction data is present beneath the slope with a thickness of 0.8 to 6 km at the base of the crust. The HVL is widely present at the base of the crust in the eastern

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portion of the northern margin of the SCS, including the outer shelf and slope. This limited distribution,

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especially its absence in the COT to the west, indicates that the HVL is not a necessary sign of the COT at the northern margin of the SCS. The magmatic intrusions and HVL may be related to partial melting caused by the decompression of a passive, upwelling asthenosphere, which resulted primarily in

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post-rifting underplating and magmatic emplacement or modification of the crust.

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3) We propose an intermediate mode of rifting to fit characteristics of the mid-northern margin of the SCS, which is different from the two passive margin end-members. This margin seems to more closely

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resemble the characteristics of magma-poor margins than those of volcanic margins.

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Acknowledgments

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This work is financially supported by the Knowledge Innovation Frontier Sciences Program of the Chinese Academy of Sciences (No.SIDSSE-201403) and the Strategic Priority Research Program of the

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Chinese Academy of Sciences (XDA11030102). We are grateful to CNOOC and SCSIOCAS for their permission to release the seismic data. Gravity modeling was performed in the Shengli Oilfield by using the FUGRO/LCT Gravity & Magnetic software of the United States. We offer many thanks to reviewer Dieter Franke and an anonymous reviewer who spent precious time to provide constructive comments that greatly improved this manuscript. References Arfai, J., Franke, D., Gaedicke, C., Lutz, R., Schnabel, M., Ladage, S., Berglar, K., Aurelio, M., Montano, J., Pellejera, N.,

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349-362.

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ACCEPTED MANUSCRIPT Chen, D.,Wu, S., Dong, D., Mi, L., Fu, S., Shi, H., 2013. Focused fluid flow in the Baiyun Sag, northern South China Sea:

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miniplate:

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http://dx.doi.org/10.1029/2007TC002216.

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TC3008,

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Figure captions

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Fig. 1 a: Geomorphology and bathymetry of the Northern South China Sea with the location of the basins. Upper left inset indicates the tectonic position. DI: Dongsha Islands, XI: Xisha Islands, ZI: Zhongsha Islands, TXB: Taixi Basin, TXNB: Taixinan Basin, PRMB: Pearl River Mouth Basin, QDNB:

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Qiongdongnan Basin, BGB: Beibu Gulf Basin, YGHB: Yinggehai Basin. Rectangle shows the location of

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study area. OBS1995 is from McIntosh et al. (2005); OBS-T3 is from Lester et al. (2014); OBS2001 is from Wang et al. (2006); ESP-E, ESP-W are from Nissen et al. (1995a); OBS2006-3 is from Wei et al.

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(2011); OBS1993 is from Yan et al. (2001); OBS2006-1 is from Ding et al. (2012), and OBH1996 is from

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Qiu et al. (2001). b: Bathymetric map and Structural units of the Pearl River Mouth Basin (PRMB) with

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adjacent sea basins. NFT, Northern Fault Terrace; ZhuID, Zhu I Depression; ZhuIID, Zhu II Depression; ZhuIIID, Zhu III Depression; DSR, Dongsha Rise; CSD, Chaoshan Depression; PYS, Panyu Swell;

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SHASR, Shenhu-Ansha Rise; SR, Southern Rise; BYS, Baiyun sag; LWS, Liwan sag.

Fig. 2 Stratigraphic column with formation, main seismic reflectors, tectonic events and unconformities of the Pearl River Mouth Basin revised from Li (1993), Wu (1994), Lüdmann and Wong (1999), Yan et al. (2001). The relative sea level curve was modified from Xu et al. (1995).

Fig. 3 Comparison of Moho depth among depth model in this study and Hu et al. (2009) and OBS model of Yan et al. (2001). Solid line represents the seismic line L2, dotted line represents the seismic line L3, and dashed line represents the seismic line L4. 41

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Fig. 4 Multi-channel seismic reflection line L1. For location of the line see Fig. 1b.

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Fig. 5 Multi-channel seismic reflection line L2. For location of the line see Fig. 1b.

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Fig. 6 Multi-channel seismic reflection line L3. For location of the line see Fig. 1b.

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Fig. 7 Multi-channel seismic reflection line L4. For location of the line see Fig. 1b.

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Fig. 8 Detailed views of the structures and reflections between the lower slope and oceanic basin in the

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seismic sections. See Figs. 5, 6 and 7 for locations.

Fig. 9 Detailed views of deep reflections. See Fig. 7 for location. Three main reflections (R1, R2 and R3)

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can be identified in upper panel, which corresponds to the top reflections of the Moho, the HVL and the lower crust, respectively. There also exists a local reflection between the upper crust and the lower crust, which is marked with „?‟ in lower panel.

Fig. 10 Two-dimensional gravity modeling along line L2. Density values are given in g/cm3. a. Step I: Density parameters of Table 1 were assigned to each layer (including seawater, sedimentary layers, crust and mantle). The long wavelength variations of calculated gravity were well correlated with the Moho depth, but there was significant misfit with observed gravity. b. Step II: We assigned a density of 2.97g/cm3 corresponding to a seismic velocity of 7-7.5 km/s and the gabbroic lithology to the HVL identified in the 42

ACCEPTED MANUSCRIPT seismic profile L2, and the fit was significantly improved, but there was still an offset between the calculated gravity and observed gravity. c. Step III: an uppermost mantle layer with a thickness of 10 km or

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so was assigned to a density of 3.0-3.15 g/cm3 in our final model, and the calculated gravity fits to the

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observed data well. d: Observed gravity, calculated gravity and error of three steps in gravity modeling.

Fig. 11 Two-dimensional gravity modeling along line L3. Density values are given in g/cm3. The steps are

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similar to those of seismic profile L2.

Fig. 12 Two-dimensional gravity model along line L4. Density values are given in g/cm3. a. Step I: Density

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parameters of Table 1 were assigned to each layer (including seawater, sedimentary layers, crust and

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mantle) which is similar to those of seismic profile L2 and L3. b. Step II: An uppermost mantle layer with

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a thickness of 10 km or so was assigned to a density of 3.0-3.15 g/cm3. Although the calculated gravity fits to the observed gravity generally, there was still an offset in Baiyun sag, where the calculated gravity was

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obvious lower than observed gravity. c. Step III: A local layer with a thickness of 2 km and a density of 2.97g/cm3 below the depression in the base of crust, and then the calculated gravity fits to the observed data well. d: Observed gravity, calculated gravity and error of three steps in gravity modeling.

Fig. 13 Detailed views of the igneous sills in the seismic sections. See Figs. 6 and 7 for locations.

Fig. 14 Comparison of the crustal structure along Profiles (a) OBS1995 (McIntosh et al., 2005), (b) OOS2004 (Xia et al., 2010), (c) OBS2001 (Wang et al., 2006), (d) ESP-E (Nissen et al., 1995a), (e) OBS2006-3 (Wei et al., 2011), (f) OBS1993 (Yan et al., 2001), (g) OBS-T3 (Lester et al., 2014), (h) 43

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northern SCS. See Fig. 1a and Fig. 15 for locations.

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Fig. 15 Distribution of the HVL in the northern SCS. White line shows the extent of the HVL. Basin and geological names see Fig.1. Grey dashed line shows the locations of previous ESP, OBH, OBS and OOS

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shows the location of multi-channel seismic lines.

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profiles, OOS2004 profiles from Xia et al. (2010), and details of other profiles see Fig. 1a. Black line

Fig. 16 Schematic three-dimensional representations of the northern margin of the SCS showing the

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lithospheric architectures (not to scale). a. Classical magma-poor margin (e.g., Iberia-Newfoundland), b.

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Classical volcanic margin (e.g., NW Europe-E Greenland), and c. intermediate margin (e.g., mid-northern

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margin of the SCS). The bathymetry and two end-members of passive margins are modified by

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Péron-Pinvidic and Manatschal (2009), Sawyer et al. (2007) and Franke (2013).

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Density (g/cm3)

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Sea water Post-rift sediments (T40-T0) Post rift sediments & Carbonates (T60-T40) Drift sediments & syn-extension sediments(T70-T60) Syn-rift sediments (Tg-T70) Continental Crust Oceanic Crust Mantle

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2.50 2.60 2.70 2.90 3.20

ACCEPTED MANUSCRIPT Highlights 1. The COT is characterized by rift depressions, structural high, and/or volcanic zone, and/or tilted fault

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blocks.

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2. The High Velocity Layer (HVL) is thickest in Dongsha uplift and thins to Manila Trench and Baiyun sag.

3. The HVL may result from post-rifting underplating and magmatic emplacement or modification of the

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4. An intermediate rifting mode is proposed to explain the mid-northern margin of the South China Sea.

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