Crustal stretching style variations in the northern margin of the South China Sea

Crustal stretching style variations in the northern margin of the South China Sea

Accepted Manuscript Crustal stretching style variations in the northern margin of the South China Sea Yongliang Bai, Dongdong Dong, Sascha Brune, Shi...

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Accepted Manuscript Crustal stretching style variations in the northern margin of the South China Sea

Yongliang Bai, Dongdong Dong, Sascha Brune, Shiguo Wu, Zhenjie Wang PII: DOI: Reference:

S0040-1951(18)30424-4 https://doi.org/10.1016/j.tecto.2018.12.012 TECTO 128003

To appear in:

Tectonophysics

Received date: Revised date: Accepted date:

14 May 2018 7 December 2018 11 December 2018

Please cite this article as: Yongliang Bai, Dongdong Dong, Sascha Brune, Shiguo Wu, Zhenjie Wang , Crustal stretching style variations in the northern margin of the South China Sea. Tecto (2018), https://doi.org/10.1016/j.tecto.2018.12.012

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ACCEPTED MANUSCRIPT Crustal Stretching Style Variations in the Northern Margin of the South China Sea

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Yongliang Baia, b*, Dongdong Dongc, Sascha Bruned, e, Shiguo Wuf, Zhenjie Wanga, b

School of Geosciences, China University of Petroleum, Qingdao 266580, China

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Laboratory for Marine Mineral Resources, Qingdao National Laboratory for Marine Science and

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a

c

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Technology, Qingdao, 266071, China

Key Laboratory of Marine Geology and Environment, Institute of Oceanology, Chinese Academy

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of Sciences, Qingdao 266071, China

Geodynamic Modelling Section, GFZ German Research Centre for Geosciences, Potsdam,

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Germany

Institute of Earth and Environmental Sciences, University of Potsdam, Potsdam, Germany

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

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e

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Abstract

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China

Linking deep seismic profiles with regional-scale gravity inversion is a powerful tool to deduce the architecture of rifted margins and their structural evolution. Here we map upper and lower crustal thicknesses of the northern South China Sea (SCS) margin in order to investigate the occurrence of depth-dependent crustal extension from the proximal to the

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Corresponding author, E-mail address: [email protected] (Y.Bai).

ACCEPTED MANUSCRIPT distal margin. By comparing upper and lower crustal stretching factors, we find that the northern margin of the SCS is segmented in three parts: (1) sedimentary basins where upper crust is stretched more than lower crust, (2) distal margin where lower crust is stretched more than upper crust, (3) mostly proximal margin regions where the two layers

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have similar stretching factors. Our results suggest that sedimentary basins and distal

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margin prominently feature depth-dependent extension, however accommodated by

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different processes. While differential thinning within sedimentary basins appears to be governed by lateral pressure variations inducing lower crustal flow, we suggest the distal

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margin to be affected by a combination of mantle flow-induced lower crustal shearing and

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sequential fault activity during crustal hyper-extension.

Keywords: crustal stretching style; lower crustal flow; the northern margin of the South China

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Sea; gravity inversion; sediment load; divergent mantle flow

ACCEPTED MANUSCRIPT 1. Introduction Several mechanisms have been suggested to describe lithospheric thinning during rifting and continental break-up. The pure shear thinning model of McKenzie (1978) assumes that each lithospheric layer at different depths exhibits the same amount of

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stretching. This kind of model has been successfully applied to many intracontinental rift

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basins where extension stopped during an early stage of rift evolution. Studies on rifted

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passive margins, however, suggest that lithospheric thinning during the later stages of rifting becomes significantly depth-dependent (Huismans and Beaumont, 2014; Kusznir and

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Karner, 2007). Classical examples of depth-dependent extension at rifted continental

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margins include the magma-poor Iberia-Newfoundland conjugates (Crosby et al., 2008; Péron-Pinvidic and Manatschal, 2009), the magma-rich Norwegian rifted margin (Roberts et

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al., 1997; Tsikalas et al., 2008), the northwest Australian margin (Baxter et al., 1999; Driscoll

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and Karner, 1998; Reeve et al., 2016) and the South China Sea margins (Clift and Lin, 2001; Franke et al., 2011; Lei et al., 2013). Depth-dependent thinning has also been suggested for

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sediment basins (Badley et al., 1988; Zhang et al., 2014).

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Different kinematic descriptions of depth-dependent extension have been suggested (Beaumont et al., 1982; Rowley and Sahagian, 1986; Royden and Keen, 1980) and several dynamic mechanisms have been proposed as an underlying reason: (1) more efficient crustal necking than mantle lithospheric thinning, which leads to exhumation of subcontinental mantle (Huismans and Beaumont, 2011; Whitmarsh et al., 2001) (2) decoupling of deformation in crust and mantle, which generates a much larger rift width at the surface than at mantle depths (Brun, 1999; Huismans and Beaumont, 2008) (3)

ACCEPTED MANUSCRIPT divergent mantle flow within the lithosphere, which generates vertical gradients in horizontal velocity (Jeanniot et al., 2016; Kusznir and Karner, 2007) (4) lateral flow of the ductile middle or lower crust, which balances upper crustal faulting and delays crustal break-up (Brune et al., 2014; Huismans and Beaumont, 2011). The individual selection for one of these

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mechanisms is thought to be controlled by a complex interplay of thermal configuration

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(Davis and Lavier, 2017), crustal thickness (Ros et al., 2017), extension rate (Tetreault and

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Buiter, 2017), rheology (Brune et al., 2017) and inherited heterogeneities (Manatschal et al., 2015). These mechanisms can even differ within a single rift between multiple phases

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(Brune et al., 2016; Naliboff et al., 2017; Svartman Dias et al., 2016). Note that in several

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natural case examples, there is ongoing controversy about the overall involvement of depth-dependent thinning (Crosby et al., 2011; Reston, 2007) and alternative processes

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have been suggested in order to explain observed structures and subsidence patterns such

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as polyphase faulting (McDermott and Reston, 2015), dynamic topography (Yang and Gurnis, 2016) or magmatic intrusion (Shi et al., 2005).

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Here we focus on the northern margin of the South China Sea (SCS) with the aim to

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quantify the role of lower crustal flow and its impact on differential crustal stretching during the evolution of a wide rift, where the width of extended crust on this margin exceeds 400 km (Hayes and Nissen, 2005; Lei et al., 2018). The SCS is one of the classical natural examples of depth-dependent extension, where the occurrence of lower crustal flow has been proposed based on several publicly available seismic profiles (Clift et al., 2002; Davis and Kusznir, 2004; Lei and Ren, 2016; Lei et al., 2015; Li et al., 2016; Zhang et al., 2008). The traditional method for estimating upper crustal stretching is based on analyzing

ACCEPTED MANUSCRIPT normal fault geometries imaged by seismic profiles. This method, however, is problematic since insufficient seismic resolution could lead to 50% unrecognized faulting within the upper crust (Clift et al., 2002; Crosby et al., 2011; Reston, 2007; Walsh and Watterson, 1991), especially if multiple fault generations exist (Reston, 2007). In this study, we employ gravity

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inversion in order to map the boundary between upper and lower crust (the Conrad interface)

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as well as the boundary between crust and mantle (Moho discontinuity) on the northern

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continental margin of the SCS. We constrain our models by available deep seismic profiles and calculate upper and lower crustal stretching factors in order to test the depth-dependent

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thinning hypothesis in our study area. Finally, we use lateral distributions of stretching

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factors in order to infer the mode of crustal stretching and the spatio-temporal occurrence of

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lower crust flow.

2. Gravity inversion methodology

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The stretching factor (  ) is defined as the ratio between initial crustal thickness ( tc0 )

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and the final present-day crustal thickness ( tc p ) (McKenzie, 1978).

  tc0 / tc p

(1)

Since the initial pre-rift thickness is not accessible, it has to be inferred from geophysical studies of the undisturbed neighboring area (Hayes and Nissen, 2005; Hu et al., 2009). The top and bottom interfaces of upper and lower crust have to be mapped in order deduce present thickness, which we do by isolating Conrad and Moho discontinuities along with bathymetry and sediment thickness data.

ACCEPTED MANUSCRIPT Free air gravity anomaly ( g faa ) in marine areas is an integration of the following components:

g faa  g mra  glra  gb  g s  go

(2)

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where g mra and glra are mantle residual anomaly and lower crust residual anomaly, which represent the gravity effects of Moho and Conrad surface undulations respectively (Figure 1).

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gb and g s are the gravity anomalies caused by the presence of seawater and sediment

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layers respectively. g o represents other components, such as density perturbations in crust

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and mantle.

For Moho or Conrad geometry inversion, we need to isolate g mra or glra by removing

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all other components from free air gravity anomaly. We assume that seawater has a constant density of 1.03 g/cm3 and so gb can be forward calculated based on bathymetry

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data. However, sediment density is depth-dependent due to load compaction. Sediment density variations depending on buried depth are modelled based on density-depth

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relationships (Sawyer, 1985), and g s is calculated based on sediment density distributions, seafloor and sediment bottom geometries. Mantle density perturbation relative to normal

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mantle affected by temperature and pressure fields are incorporated by means of thermal expansion coefficient (Chappell and Kusznir, 2008; Kusznir et al., 2018), pressure-driven compressibility coefficient (Bouhifd et al., 1996; Kroll et al., 2012), temperature modelling (McKenzie, 1978) and pressure modelling (Afonso et al., 2008). Note that the gravity effect of this mantle density perturbation is the main part of g o and we ignore all the other components of free air gravity anomaly. For details about our modelling methods and parameter selections, we refer to a previous article (Bai et al., 2014).

ACCEPTED MANUSCRIPT The workflow for Conrad and Moho interface inversion is shown in Figure 2. There are two key input parameters for gravity inversion across density contrasts based on target residual anomaly, which are density contrast value and average buried depth (also called reference depth) of the interface (Cowie and Kusznir, 2012; Oldenburg, 1974; Parker, 1973).

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We use discretized density ranges for upper crust and lower crust within a specific interval,

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and different possible reference depths for Moho and Conrad. All possible parameters

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combinations are tested in the inversion flow (Figure 2). The workflow minimizes for smallest root mean square (RMS) differences with seismic interpretation results in order to provide

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the most suitable values for density contrast and reference depth. We set the possible value

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range for the density of continental upper and lower crust between 2.5 g/cm3 to 3.1 g/cm3 (Gao et al., 2015). Further we define possible depth ranges for the Moho (between 20 km to

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30 km), and for the Conrad interfaces (between 10 km to 20 km).

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3. Geological background of the SCS

The formation of the magma-poor northern continental margin of the SCS is affected by

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two independent geologic process: the lateral extrusion of Indochina during India-Eurasia collision (Briais et al., 1993; Hao et al., 2011; Lee and Lawver, 1994; Mazur et al., 2012), and the proto-SCS southward subduction (Hutchison et al., 2000; Zahirovic et al., 2014), while it is also possible that these two mechanisms acted together (Cullen et al., 2010; Rangin et al., 1999). The submerged South China Block covered most of the thinned continental region and thinning onshore is limited (Hayes and Nissen, 2005; Hu et al., 2009). There are at least

ACCEPTED MANUSCRIPT two extension stages according to subsidence and fault analysis, one is from the Late Cretaceous to Eocene with NW-SE extension; and the other one is from Late Eocene to Early Miocene involving N-S extension (Lee and Lawver, 1994). On the northern margin of the SCS, several large basins with thick sediments formed during long-term extensional and

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thermal-subsidence phases. The Pearl River Mouth Basin is the largest depocenter within

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our study region and experienced anomalously high post-rift subsidence rates after the

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initiation and subsequent cessation of seafloor spreading (Xie et al., 2014). There is a controversy about the lithospheric stretching style of the northern margin of

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the SCS due to different analyses of stretching factors of upper crust, crust and lithosphere.

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Su et al. (1989) suggest a depth-uniform stretching style for the Pearl River Mouth Basin. Later studies however highlighted depth-dependent thinning of the lithosphere by estimating

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upper crustal stretching factor based on normal fault geometries, full crustal stretching factor

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from seismic velocity interface interpretations and lithosphere-scale stretching factor from sediment subsidence analysis (Clift et al., 2002; Davis and Kusznir, 2004; Lü et al., 2017;

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Lei and Ren, 2016; Lei et al., 2015; Li et al., 2016; Zhang et al., 2008). This kind of

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subsidence analysis however is based on the assumption of instantaneous stretching, which does not consider heat loss during rifting. This contrasts the inference that basin formation is a consequence of horizontal extension and vertical thermal subsidence (Badley et al., 1988; Sclater and Christie, 1980). Chen (2014) designed a method that considers both structural and geothermal subsidence stages and applied it to the proximal part of the Pearl River Mouth Basin north of the Baiyun sub-basin and found a depth-uniform stretching style suggesting that former studies overestimated the lithospheric stretching factor. Here we use

ACCEPTED MANUSCRIPT gravity inversion and deep seismic profiles in order to provide new insight into the crustal stretching style at the northern margin of the SCS.

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4. Employed data sets We build on several geophysical data sets that have been collected for scientific studies

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and hydrocarbon explorations. There are 9 publicly available deep seismic profiles (Figure

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3), ESP1985-2 (Hayes and Nissen, 2005), OBS1993 (Yan et al., 2001), OBH1996-4 (Qiu et

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al., 2001), OBS2001 (Wang et al., 2006), OBS2006-1 (Li et al., 2011), OBS2006-2 (Ao et al., 2012), OBS2006-3 (Wei et al., 2011), OBS2011 (Huang et al., 2011), OBS2013-3 (Guo et al.,

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2016). All of these profiles provide Moho and Conrad interpretation results except for

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ESP1985-2, which does not specify the Conrad interface. The seismic velocity of the Conrad

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interface for profile OBS2001 by Wang et al. (2006) is defined to be 6.5 km/s, for OBS2006-3 by Wei et al. (2011) is 6.8 km/s, and for all the other six profiles it is 6.4 km/s.

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Therefore in this study, we use 6.4 km/s as the seismic wave velocity definition of the Conrad interface. Since the Conrad interface represents the boundary between upper crust

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and lower crust, it also provides a density contrast (Berckhemer, 1969; Wang et al., 2007). We exploit this fact by constraining our gravity inversion with the six profiles mentioned above (Figure 3 and Figure 4). Even though seismic data provides relatively accurate interface perturbations, uncertainties arise during data acquisition, processing and interpretation. For example, the discrepancy of Conrad depth from OBS2011 and OBS2013-3 on the intersection point

ACCEPTED MANUSCRIPT between them is 3.8 km, and the Moho depth discrepancy is 1.6 km; the discrepancy of Conrad depth from OBS2006-2 and OBH1996-4 on their intersection point is 4.3 km, and the Moho depth discrepancy is 3.5 km. Considering seismic profiles are the most accurate reference data available in such a large coverage in our study region, it is reasonable to take

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seismic interpretations as the constrain and standard for gravity study.

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We use free-air gravity anomaly from the open database of Sandwell et al. (2014).

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Bathymetry and sediment thickness data, which are used to remove seawater and sediment gravity effect is extracted from ETOPO1 (Amante and Eakins, 2009) and the NGDC global

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sediment thickness grid (Divins, 2004), respectively. For the oceanic crustal age grid which

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is necessary for thermal modelling we use magnetic lineation interpretations by Müller et al. (2008). Upper and lower crust thickness uncertainty is affected by the uncertainty of

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bathymetry and sediment thickness data, also the uncertainty of our gravity interface

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inversion result (Kusznir et al., 2018). The RMS between the depth of the sediment basement from seismic interpretation and the NGDC sediment thickness and ETOPO1

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bathymetry and topography grid along the profiles where the sediment basement was

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interpreted is 1.1 km. This discrepancy mainly locates in sediment basins, especially with thick sedimentations.

ACCEPTED MANUSCRIPT 5. Results 5.1 Crustal thickness Preferable Moho and Conrad inversion result is derived based on the parameters shown

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in Table 1 through comparing root mean square (RMS) difference between seismic

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interpretations and gravity inversion results based on each parameter combination (density contrast and reference depth). RMS between Moho/Conrad and seismic interpretations is

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1.9/2.2 km respectively. The RMS for Moho inversion result here is reduced compared with

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our former studies in the SCS conjugate margins based on similar methodology (Bai et al.,

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2014; Bai et al., 2015), 1.9 km VS 2.7 km. The difference is that we ignored the density contrast between oceanic and continental crust, and between upper and lower crust in our

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former studies. In this study, however, we set the density of oceanic crust to be 0.15 g/cm3

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larger than continental crust, whereas the necessary COB line is based on Bai et al. (2015). We benchmark our crustal thicknesses that are derived from gravity inversion to six

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deep seismic profiles in Figure 4. The larger structures of the seismic profiles are very well

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reproduced with our inversion technique. Note however, that we cannot resolve small wavelength structures since the seismic resolution is higher than the gravity grid. According to the comparison of upper and lower crustal thickness data from different sources (Figure 5b, c and Figure 4), relatively large discrepancies of the upper crustal thickness occur when there are considerable sediment thickness discrepancies from NGDC and seismic profiles. The profile OBS2013-3 provides thicker lower crust and thinner upper crust than OBS2011 (Figure 4), and also provides larger upper and lower crustal thickness differences from

ACCEPTED MANUSCRIPT gravity and seismic (Figure 5b and c). On the whole, the RMS between upper/lower crust thickness by gravity inversion and seismic interpretation is 2.4 km and 2.5 km respectively. Further uncertainties arise due to the fact that data sources are not necessarily consistent. For example, the sediment thickness on northwest end of the profile OBS2011 by Huang et

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al. (2011) is ~3 km, but the NDGC sediment thickness here is 7 km (Figure 4, Figure 5a).

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This uncertainty has to be kept in mind when interpreting our results.

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We further find that the top basement bends downward in the regions with thick sedimentation, for example the Pearl River Mouth Basin (PRMB) along OBS2006-1, Xisha

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Trough along OBH1996-4 and OBS2011, ZhongJinanNan Basin (ZJNB) along OBS2013-3

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(the pink-rectangle bounded regions in Figure 4). Conrad has similar trend with Moho, both of them lack short wavelength undulations independent of sediment thickness variations

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(Figure 4). So, lower crustal thickness variations are much smoother than those of the upper

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crust since the later ones highly depend on short-wavelength variations of sediment thickness (solid line curves in Figure 4). Upper and lower crustal thickness based on gravity

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inversion display similar variations in these regions (dashed line curves in Figure 4).

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We present our inversions for upper, lower and total (excluding sediment) continental crustal thicknesses in Figure 6. Our total crustal thickness has similar trends with the gravity inversion result by Gozzard et al. (2018) and we also provided quantitative uncertainty tests by taking deep seismic profiles as reference (Figure 4 and Table 1). In sedimentary basins, we find that the upper crustal thickness has a strong negative correlation with sediment thickness, while the lower crustal thickness distribution shows a much weaker inverse correlation. In the distal margin, the lower crust is much thinner than the upper crust. Outside

ACCEPTED MANUSCRIPT of these specific domains, our research area shows similar upper and lower crustal thickness distributions. 5.2 Stretching factor

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For the calculation of stretching factors we follow Hayes and Nissen (2005) in assuming

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that the initial unstretched crustal thickness equals to average crustal thickness of SE China . There are different initial thickness estimates for the northern continental margin of the SCS

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based on geophysical studies: 32 km from Hu et al. (2009), 35 km by Hayes and Nissen

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(2005), ~30–35 km by Li et al. (2006). In the following, we use the intermediate value of 32

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km as our reference value, but we perform additional robustness tests by accounting for end-member cases in the next paragraph. We compute stretching factor maps for upper and

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lower crust (Figure 7a and Figure 7b) by assuming that the initial upper and lower crustal

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thickness was 15.1 km and 16.9 km respectively. These values represent the average thicknesses along the present-day contour of our reference unstretched full-crustal

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thickness of 32 km.

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Our result show key trends for the upper and lower crustal stretching factors. In general, both stretching factor distributions increase from onshore towards oceanic basins, but they are also increasing with sediment thickness in proximity to the major basins (Figure 7). Especially upper crustal stretching appears to correlate with sediment thickness, where peak values reach more than 4 in the Qiongdongnan and about 3 in the Taixinan Basin, respectively. In contrast, lower crustal stretching factors are highest near the continent ocean boundary.

ACCEPTED MANUSCRIPT We focus on the relative impact of upper versus lower crustal stretching by plotting their difference in Figure 8. This map clearly illustrates that the continental crust did not stretch in a constant way but that it involves a distinct component of depth-dependency. We divide the resulting map in three regions that correspond to three distinct styles of stretching: (1) in

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sedimentary basins, the upper crust has been stretched more than the lower crust and the

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stretching factor difference is correlated with sediment thickness; (2) in the distal margin, the

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lower crust has been stretched more than the upper crust; (3) in areas outside of these regions, which includes the main part of the continental northern margin of the SCS, upper

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and lower crust have similar stretching factors. The underlying processes that might be

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responsible for this pattern are discussed in the following section. In order to test the impact of the assumed initial upper and lower crustal thickness, we

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employ two end-member values that are listed in Table 2. We find that within the possible

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initial thickness range, different initial values lead to similar stretching factor distributions (Figure 7). Especially in the regions with thick crust, sedimentary basins and distal margin,

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upper and lower crustal stretching is more pronounced for the thicker initial crust.

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Nevertheless, the main trend of these three different stretching factor map groups are quite similar, corroborating the robustness of our conclusions.

6. Mechanisms for basin subsidence, crustal thinning, and lower crust flow In general, basin subsidence and the associated creation of accommodation space can result from isostatic adjustment due to several processes (Ingersoll, 1988): (1) crustal

ACCEPTED MANUSCRIPT thinning in extensional environment; (2) sediment loading and compaction; (3) lateral flow of ductile crust away from the basin axis and (4) thermal post-rift subsidence due to lithospheric cooling. Usually, these factors do not affect the subsidence process independently but they interact mutually. For example, normal faulting in the brittle upper

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crust is significantly affected by sediment loading (Buck, 1988), while sediment loading may

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be one of the driving forces for lower crust flow (Clift et al., 2015; Westaway, 2002). Of these

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four processes only the first three can lead to crustal-scale depth-dependent thinning. We suggest that all of them contributed to shaping the northern margin of the SCS, but that their

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impact varies within specific margin domains.

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6.1 Sedimentary basins

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Figure 9 illustrates two potential causes of lower crustal flow in sedimentary basins that

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might occur in sequential order. During the first stages of extension, it can be expected that the whole crust is thinned and extended as a whole where deformation is accommodated by

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normal faulting and ductile shear zones. With ongoing crustal and lithospheric thinning, mantle upwelling takes place beneath the basin axis, which increases the crustal

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temperature and hence reduces its viscosity. In this case, compensation or at least partial compensation takes place in the weak lower crust (Brune and Ellis, 1997; Buck, 1991; Hopper and Buck, 1998; Voorde et al., 1998). When the sediment layer is thin (Figure 9a), the lithostatic pressure at the Conrad interface increases away from the basin center. This is illustrated in Figure 9c, where the black dashed curve shows variations of PRMB loading pressure in a hypothetical situation without sediments. In this case, there is an obvious

ACCEPTED MANUSCRIPT pressure deficit in the basin, especially when the starved basin is deep enough. Hence, in this first stage, flow of the low-viscosity lower crust is expected to occur from high to low pressure towards the basin axis (Bertotti et al., 2000; McKenzie et al., 2000; Morley and Westaway, 2006; Zhao et al., 2018). This mechanism is similar to the suggested lower

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crustal flow of the Tibetan plateau as both processes are driven by pressure differences

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from regions of thick crust to areas of thin crust (Clark and Royden, 2000; Royden et al.,

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2008). Note however, that lower crustal flow in the SCS acts on much shorter lengths scales. If there is abundant sedimentary supply allowing the basin to be filled (Figure 9b), the

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loading pressure beneath the basin increases. The black solid curve in Figure 9c shows the

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current loading pressure on the Conrad interface in the PRMB and the red curve shows the pressure increase by sediment loading. Beneath the basin axis, we find that the pressure

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has been increased by up to ~70 MPa. The pressure increase could induce lower crustal

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outward flow away from basin axis (Figure 9b) (Clift et al., 2015; Morley and Westaway, 2006). Note that the present-day configuration in Figure 9c represents the equilibrated

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situation, which for high sedimentation rates may have been preceded by a situation where

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the loading pressure within the basin exceeded the one outside of the basin inducing lower crustal flow outward from the basin axis (Figure 9b) (Clift et al., 2015; Morley and Westaway, 2006). High sedimentation rates that are required to drive such a process have been inferred at the mouths of major rivers draining into the SCS (Clift and Sun, 2006) but also other South East Asian rivers (Brune et al., 2016; Clift and Sun, 2006), which transport large amounts of material from erosional areas of the Tibetan plateau. As an additional effect, thermal blanketing due to thick sediment coverage increases the temperature within the crust and

ACCEPTED MANUSCRIPT raises the brittle-ductile transition bringing parts of the upper crust into the ductile regime (Bertotti et al., 2000; Sun et al., 2008). Hence the bottom part of upper crust would flow together with the lower crust leading to additional upper crustal thinning (Figure 9b). Sediment maximum density is 2.64 g/cm3 based on the sediment density modelling

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method (Sawyer, 1985) and the parameters by Clift et al. (2002) based on ocean drilling

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program (ODP) sites on the northern margin of the SCS, which is similar to upper crust

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density (2.66 g/cm3, Table 1). When the buried depth gets shallower, the density contrast between sediment and upper crust becomes larger; which means average sediment density

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is smaller than upper crust. So no matter the basin has been filled up or not, pressure on

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lower crust is deficit in basin, especially in basin axis, relative to basin surroundings. For example, Figure 9c shows that the present-day PRMB still exhibits a deficit of pressure at

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the Conrad interface. And this pressure deficit is positively correlated with sediment

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thickness. So on the whole, despite outward lower crustal flow due to the process illustrates in Figure 9b, the pressure deficit within the basin would generate a net inward lower crust

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the basins.

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flow, which is why the volume of lower crust correlates positively with sediment thickness in

In combination, the described processes lead to more pronounced thinning of the upper crust than of the lower crust. This explains why sediment thickness correlates with predominantly upper crustal thinning (Figure 8). The main difference between our findings and former studies is that the crustal stretching style appears to be neither uniform (Chen, 2014) nor that the stretching factor increases with depth (Davis and Kusznir, 2004; Zhang et al., 2008) in the main sedimentary basins, but that instead the stretching factor of the crust

ACCEPTED MANUSCRIPT decreases with depth. This finding is further corroborated by deep seismic profiles (Figure 4) and previous interpretations in the Baiyun sub-basin, PRMB and Taixinan Basin (McIntosh

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et al., 2014; Sun et al., 2008).

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6.2 Distal margin

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According to our inversion results, lower crust has been stretched significantly in the SCS distal margin (Figure 7), with a stretching factor that exceeds the one of the upper crust

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(Figure 8), and that is also larger than the lower crustal stretching factor of the proximal

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margin. These results can be explained by two independent processes discussed in more detail below: (1) upwelling mantle exerting shear stresses at the base of the crust (2)

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basin-ward localization and crustal hyper-extension.

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Lithospheric necking always results in mantle upwelling. Numerical models indicate that the low-viscosity asthenosphere rises from deep under the rift center and diverges at

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shallow depths (Geoffroy et al., 2015; Huismans and Beaumont, 2011). This process takes

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place during passive rifting where the far-field forces drive extension (Brune, 2018), but is even more pronounced in situations where the asthenospheric buoyancy contributes an additional driving force (Mondy et al., 2018). If the mantle lithosphere is split prior to crustal break-up (Huismans and Beaumont, 2011), which has been recently inferred for the SCS (Larsen et al., 2018) the divergent asthenospheric flow field may transmit continent-ward directed tractions directly to the lower crust. Thereby it may drag the lowermost crust away from the rift center and hence lead to differential crustal thinning of the lower crust (Figure

ACCEPTED MANUSCRIPT 10a). This effect will be more pronounced if the local extension rates increase during rift evolution as indicated by plate reconstructions and numerical modelling (Brune et al., 2016; Brune et al., 2018), as well as kinematic block reconstructions at rifted margins (Sutra and Manatschal, 2012). The continent-ward directed tractions can be transmitted to the lower

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crust and are expected to result in continent-ward dipping normal faults that have been

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interpreted in some seismic cross section of the SCS (Clerc et al., 2018).

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A second effect that leads to more pronounced lower crustal thinning near the distal margin is basin-ward localization. During the final stages of rifting, the crust has been highly

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thinned potentially leading to embrittlement of the lower crust. Hence, faults can cut through

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the whole crust and reach the mantle, and thus the faults become more effective for whole crustal thinning (Pérez‐Gussinyé and Reston, 2001). Prominent brittle deformation favors

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lower crustal thinning due to successive fault activity (Figure 10b) (Brune et al., 2017;

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Ranero and Pérez-Gussinyé, 2010) which finally leads to lower crust being stretched

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significantly more than upper crust in the distal margin. This process explains the extreme thinning in the lower crust, which can be inferred for

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almost the entire distal margin domain of our study area (Figure 7,8), whereas the only exception is the Zhongsha Islands area. Here, the upper and lower crust was stretched much less than at its surroundings and these two layers here have similar stretching factors, probably because the Zhongsha Islands are rigid (Li, 2011).

ACCEPTED MANUSCRIPT 7. Conclusions We deduced Conrad and Moho geometries via gravity inversion techniques, by accounting for thermal expansion of the mantle lithosphere as well as pressure compression effects within the sedimentation layer and the mantle. Comparison to seismic interpretations

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shows that our inversion very well reproduces the long-wavelength signal of both Moho and

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Conrad interface. Stretching factors for upper and lower crust have been mapped according

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to possible initial crustal thickness values.

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Our results show that the crustal stretching styles in the northern margin of the SCS can be classified into three types representing three specific regions: (1) in sedimentary basins,

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upper crust is stretched more than lower crust, (2) the opposite case occurred in the distal margin, while (3) other areas of the continental margin appear to be uniformly stretched.

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Even though sedimentary basins and distal margin have different crustal stretching style, we infer that depth-dependent crustal thinning took place in both of these areas, albeit

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accommodated by different processes. We suggest that in sedimentary basins, the dominant process enabling crustal flow is a

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pressure gradient in the lower crust: as long as the sedimentation in basins is thin, the pressure deficit induces lower crustal flow into the basin, thus balancing upper crustal faulting with lower crustal flow, which is ultimately expressed in the higher differential thinning of the upper crust. During a later stage, however, enhanced sediment supply may increase pressure and temperature within the crust, which will result in outward flow of lower crust and lowermost parts of the upper crust. On the whole, our results suggest that the

ACCEPTED MANUSCRIPT inflow of lower crust is the dominant process explaining the relatively high differential upper crustal thinning in sedimentary basins. Depth-dependent crustal thinning near the distal margin can be caused by processes acting during the final stages of rifting, where shear stresses of the diverging mantle can be

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transmitted directly to the lower crust leading to more pronounced lower crustal thinning.

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Furthermore, the outermost part of the distal margin should be affected by crustal

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embrittlement and subsequent hyperextension, which causes further differential crustal

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thinning where the lower crust is stretched more than the upper crust. Acknowledgements

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We would like to express our gratitude to Editor Ramon Carbonell, Dr. Lei Chao and one

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anomalous reviewer for their detailed and constructive revision of our manuscript. This work

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is supported by the National Natural Science Foundation of China (No. 41506055, 41476046, 41476042, U170120019, 41506085), Fundamental Research Funds for the Central

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Universities China (No.17CX02003A, 18CX02064A), Open Fund of the Key Laboratory of Marine Geology and Environment, CAS (No. MGE2017KG01), Key Research Program of

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the Chinese Academy of Sciences (Grant No.Y10131) and the Helmholtz Association through grant number VH-NG-1132 for the Helmholtz Young Investigators Group CRYSTALS.

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Figures

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Figure 1. Cartoons show meaning of two residual gravity anomalies. (a) After removing the

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gravity effects of seawater and sediment (layers indicated by solid-line grids), the density contrast with upper crust, as well as mantle (layers covered by dashed-line grids), the

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density contrast with lower crust from free air gravity anomaly, we obtain lower crust residual

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anomaly ( glra ), which represents the gravitational effect of the Conrad interface. (b) Removing gravity effects of seawater, sediment and upper crust (the layers indicated by

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solid-line grids), the density contrast with lower crust from free air gravity anomaly allows us

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to deduce the mantle residual anomaly ( g mra ), that represents the gravity effect of the Moho

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interface.

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Figure 2. Inversion workflow for Conrad and Moho interface after Chappell and Kusznir

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(2008) and Bai et al. (2014). Based on sediment thickness (NGDC), bathymetry (ETOPO1) and Moho (CRUST1.0) data, we calculate the gravity effect of each layer and separate the

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lower crust residual anomaly ( glra ). Conrad geometry is inverted based on glra constrained by deep seismic profiles. We then use the Conrad interface to isolate the mantle residual

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anomaly ( g mra ) for Moho inversion. This Moho geometry features a higher resolution than CRUST1.0 data, which is why we use it to refine our Conrad interface inversion.

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Figure 3. Tectonic setting of the South China Sea northern margin after Bai et al. (2015) and Lei and Ren (2016). Base map is the bathymetry and elevation grid from ETOPO1 (Amante

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and Eakins, 2009). The white rectangle shows our research area including deep seismic profiles. RRF, Red River Fault; ZI, Zhongsha Islands; DR, Dongsha Rise; ZU, Zhongsha

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Uplift; TB, Taixinan Basin; PRMB, Pearl River Mouth Basin; LB, Leidong Basin; BGB, Beibu

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Gulf Basin; YB, Yinggehai Basin; QB, Qiongdongnan Basin; ZJNB, ZhongJianNan Basin.

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Figure 4. Comparison of upper crustal thickness (UCT) and lower crustal thickness (LCT)

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curves from seismic interpretation and gravity inversion; the tectonic structures are from deep seismic profiles, OBS2006-2 is by Ao et al. (2012), OBS1993 is by Yan et al. (2001),OBS2006-1 is by Li (2011), OBS2011 is by (Huang et al., 2011), OBH1996 is by (Qiu et al., 2001) and OBS2013-3 is by Guo et al. (2016). The red lines show the intersection point between the profile OBS2011 (e) and OBS2013-3 (k), and the intersection point between the profile OBS2006-2 and OHB1996-4. Profile locations are shown in Figure 3. PRMB, Pearl River Mouth Basin; XT, Xisha Trough; ZJNB, ZhongJianNan Basin.

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Figure 5. Density-contrast interface geometry or layer thickness comparisons from different

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sources. Sediment bottom buried depth from seismic VS from NGDC (Divins, 2004) and

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ETOPO1 (Amante and Eakins, 2009) (a), the RMS between them is 1.1 km. Upper and

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lower crustal thickness from seismic interpretations and gravity inversions (b and c), with RMS values 2.4 km and 2.5 km respectively. The thickness values in the red ellipses are

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along the profile OBS2013-3.

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ACCEPTED MANUSCRIPT Figure 6. Upper (a), lower (b) and total (c, excluding sediment) continental crustal thickness according to our Moho and Conrad interface inversion workflow and basement top geometry based on topography (Amante and Eakins, 2009) and sediment thickness (Divins, 2004) data. Black contours represent sediment thickness in kilometers from the NGDC (Divins,

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2004) grid. ZI, Zhongsha Islands; DR, Dongsha Rise; ZU, Zhongsha Uplift; XT, Xisha

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Trough; NWSB, North West Sub Basin; TB, Taixinan Basin; PRMB, Pearl River Mouth Basin;

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LB, Leidong Basin; BGB, Beibu Gulf Basin; QB, Qiongdongnan Basin; ZJNB, ZhongJianNan

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Figure 7. Upper and lower crustal stretching factor maps based on different assumed initial

ACCEPTED MANUSCRIPT total crustal thicknesses (AITCT) for upper and lower crust. Assumed initial upper crustal thickness (AIUCT) is 15.1 km, 14.2 km and 16.5 km for (a), (c) and (e) respectively; assumed initial lower crustal thickness (AILCT) is 16.9 km, 15.8 km and 18.5 km for (b), (d) and (f), respectively. Labeled contours represent sediment thickness in kilometers from

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NGDC (Divins, 2004). Our reference model with 32 km initial crustal thickness is shown in

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panel (a) and (b). ZI, Zhongsha Islands; ZU, Zhongsha Uplift; XT, Xisha Trough; NWSB,

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North West Sub Basin; TB, Taixinan Basin; PRMB, Pearl River Mouth Basin; LB, Leidong

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Basin; BGB, Beibu Gulf Basin; QB, Qiongdongnan Basin; ZJNB, ZhongJianNan Basin.

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Figure 8. Stretching factor differences (  ) distribution map.  is the result of upper

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crustal stretching factor minus lower crustal stretching factor. Redish domains are dominated by upper crustal stretching, while in greenish areas lower crustal stretching

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prevails. In yellow regions, upper and lower crust stretched similarly. ZI, Zhongsha Islands; DR, Dongsha Rise; ZU, Zhongsha Uplift; NWSB, North West Sub Basin; TB, Taixinan Basin; PRMB, Pearl River Mouth Basin; LB, Leidong Basin; BGB, Beibu Gulf Basin; QB, Qiongdongnan Basin; ZJNB, ZhongJianNan Basin.

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Figure 9. Mechanisms of lower crustal flow in sediment basin. (a) If sediment thickness is

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thin during ongoing extension, lower crust flows along pressure gradients from regions of thick upper crust to regions of thin upper crust. (b) Pronounced sediment loading during and

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after extension increases the lithostatic pressure beneath the basin and lower crustal flow occurs away from the basin axis. Thick sedimentation cover also raises the brittle-ductile

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transition (BDT) by means of thermal blanketing, thereby allowing for additional removal of the lowermost parts of the upper crust. (c,d) Present geologic structure of the PRMB, loading

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pressure on Conrad before and after accounting for the weight of the sediments, and pressure increase by the sediment filling. The pressure deficit beneath the basin reduces significantly due to the sediment fill. Figure 8 shows locations of this profile.

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Figure 10. Mechanisms promoting depth-dependent extension in distal margin. (a) Divergent

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mantle flow shears the lower crust. This process occurs after mantle lithospheric separation but before crustal break-up and hence affects only the distal margin. (b) Basin-ward

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localization and lower crustal embrittlement promote sequential fault activity where new

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faults form on the basin-ward side within the hanging wall of old faults, and also move the brittle-ductile transition (BDT) interface beneath the lower crust in the distal margin. This

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process leads to efficient removal of lower crust during the final stages of crustal break-up.

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Table 1. Parameter values for Moho and Conrad inversion.

upper

Oceanic crust density (g/cm3)

lower

upper

lower

Moho 2.90

2.81

3.05

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2.66

Mantle Reference RMS density depth (g/cm3) (km) (km)

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Inversion target

Continental crust density (g/cm3)

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Total crustal thickness

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32

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Upper crustal thickness

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15.1

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Graphical abstract

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1. Regional stretching factor contrasts between the upper and lower crust are mapped in the northern margin of the SCS. 2. Pressure variations due to lateral crustal thickness differences and sedimentation are the

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main mechanisms for depth-dependent stretching in the sedimentary basins 3. Mantle upwelling, basin-ward localization and crustal hyper-extension are the main

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mechanisms for depth-depend stretching at distal margin.