Crustal structure north of the Taiping Island (Itu Aba Island), southern margin of the South China Sea

Accepted Manuscript Crustal structure north of the Itu Aba Island (Taiping Island), southern margin of the South China Sea Jih-Hsin Chang, Hsien-Hsiang Hsieh, Arif Mirza, Sung-Ping Chang, Ho-Han Hsu, Char-Shine Liu, Chih-Chieh Su, Shye-Donq Chiu, Yu-Fang Ma, Ying-Hui Chiu, Hau-Ting Hung, Yen-Chun Lin, Chien-Hsuan Chiu PII: DOI: Reference:

S1367-9120(16)30258-9 http://dx.doi.org/10.1016/j.jseaes.2016.08.005 JAES 2782

To appear in:

Journal of Asian Earth Sciences

Received Date: Revised Date: Accepted Date:

18 March 2016 6 August 2016 6 August 2016

Please cite this article as: Chang, J-H., Hsieh, H-H., Mirza, A., Chang, S-P., Hsu, H-H., Liu, C-S., Su, C-C., Chiu, S-D., Ma, Y-F., Chiu, Y-H., Hung, H-T., Lin, Y-C., Chiu, C-H., Crustal structure north of the Itu Aba Island (Taiping Island), southern margin of the South China Sea, Journal of Asian Earth Sciences (2016), doi: http://dx.doi.org/ 10.1016/j.jseaes.2016.08.005

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Crustal structure north of the Itu Aba Island (Taiping Island), southern margin of the South China Sea

Jih-Hsin Chang1*, Hsien-Hsiang Hsieh1, Arif Mirza1, Sung-Ping Chang1, Ho-Han Hsu1,2, Char-Shine Liu1, Chih-Chieh Su1, Shye-Donq Chiu1, Yu-Fang Ma3, Ying-Hui Chiu1, Hau-Ting Hung4, Yen-Chun Lin5, Chien-Hsuan Chiu6

1 Institute of Oceanography, National Taiwan University, Taipei, Taiwan. 2 National Oceanic Centre, Southampton, UK. 3 Precision Instrumentation Center, National Taiwan University, Taipei, Taiwan. 4 Offshore Exploration & Production Division, Exploration & Production Business Division, Chinese Petroleum Company, Taipei, Taiwan. 5 GeoResource Research Center, National Cheng Kung University, Tainan, Taiwan. 6 Bureau of Mines, Ministry of Economic Affair, Taipei, Taiwan. *Corresponding author ([email protected]; +886-2-3366-1871)

Abstract Based on the multi-channel seismic (MCS) and gravity data offshore north of Itu Aba Island (Taiping Island) in the Spratly Islands (Nansha Islands), we revisited the crustal structures in the northern part of the southern margin of the Southwest (SW) Sub-basin of the South China Sea (SCS). The MCS data suggest that the basement structural highs in the southwest margin of the SCS are dominated by both fault blocks and volcanic basement structures that probably formed along with the basement faults. The gravity modeling results reveal that these volcanic basement

structures were probably associated with the high velocity or density lower crust (HVDLC), suggesting that the southern margin of the SW Sub-basin serves as an intermediate margin. Based on the tectonic features in the SCS margins, including (1) the transition between the magma-poor margins of the Northwest (NW) Sub-basin and intermediate margins of the East and SW Sub-basins; (2) the cessation of the seafloor spreading in the NW Sub-basin being succeeded by a major ridge jump event, which was very likely to reflect an active mantle upwelling event; (3) the extensive distribution of the HVDLC beneath in the well-investigated northern margin of the East Sub-basin and as explored in the SW Sub-basin in this study; and (4) basement faulting and volcanism shown in this study that may be related to the HVDLC and hyper-extension, we proposed a conceptual model to interpret the origin of the HVDLC in the southern margin of the SW Sub-basin. The margin of the SCS basin was magma-poor at the time it occurred. An active mantle upwelling event caused by small-scale mantle convection, possibly related to the influence of the nascent Hainan mantle plume, was formed subsequently, resulting in a southward ridge jump and the HVDLC beneath the current margins of the East and SW Sub-basins. The East and SW Sub-basins continued to spread, latitudinally dividing their HVDLCs as the margins separated accordingly. Afterwards, the hyper-extension in the distal margins may also be responsible for a crust weakening zone wherein the subsequent volcanic events could occur.

Keywords: Itu aba Island (Taiping Island), South China Sea, Hainan mantle plume, small-scale convection, hyper-extension, high velocity layer, lower crust material.

1. Introduction Passive margins can be identified as magma-poor, volcanic, and intermediate (Clift et al., 2001; Zhu et al., 2012; Gao et al., 2015; Franke et al., 2011; Doré and Lundin, 2015) (Fig. 1). The magma-poor or non-volcanic margins, typically considered in the offshore Iberian Peninsula, feature a highly stretched deformation of continental crust (Whitmarsh et al., 2001). The volcanic margins, as can be seen off the mid-Norwegian, northern North Atlantic coast, are characterized by seaward dipping reflectors (SDR) in the multichannel seismic (MCS) profiles (Mutter et al., 1982). Another important feature of the volcanic margins is the lower crust material (LCM), generally recognized as high velocity layers (HVL) or high velocity lower crust (HVLC) in the continent-ocean transition (COT). In this study, we inclusively termed them high velocity or density lower crust (HVDLC). The intermediate margins contain characteristics of both end-member types. For instance, some parts of the margins of the South China Sea (SCS) (Fig. 2) are characterized by volcanic intrusions in the MCS profiles and the HVDLC in ocean bottom seismometer (OBS) data and gravity anomaly modeling results.

The SCS oceanic basin is composed of the Northwest (NW) Sub-basin, Southwest (SW) Sub-basin, and East Sub-basin. The margins of the SCS oceanic basin were earlier considered as non-volcanic margins (Yan et al., 2001) and recently considered to be characterized as an intermediate margin (Zhu et al., 2012; Gao et al., 2015). In fact, the margin types of the SCS Sub-basins are neither identical nor fully understood. The margin of the NW Sub-basin, which was the earliest portion of the SCS opening, is magma-poor (Ding et al., 2012). The succeeding East Sub-basin is characterized by margins with large volumes of volcanic intrusions and the HVDLC beneath its

northern margin, suggesting that it features an intermediate margin (Zhu et al., 2012; Gao et al., 2015). The nature of the current SW Sub-basin, however, is still a question under debate, owing to inconsistent survey analyses and results.

The nature of the SW Sub-basin is receiving increasing attention and has recently become controversial. The feature of the northern margin of the SW Sub-basin was earlier investigated by Lü et al. (2011), concluding that the two end-member continental rifting models could not fully explain the crustal structures of the northern margin of the SW Sub-basin. Based on the OBS investigation, Qiu et al. (2011) suggested that there is no obvious HVDLC found by an OBS profile (Line NH973-1, see Fig. 2 for the location) along the dip of the conjugated southern margin of the SW Sub-basin. Nevertheless, a later OBS study along another dip profile (Line PR2, see Fig. 2 for the location) of the southern margin of the SW Sub-basin suggested that there is a HVDLC existing under the crust (Pichot et al., 2014). It is questionable if the nature of the lower crust varies that drastically in less than 200 km. More investigation and analysis are required to better understand the nature of the SW Sub-basin.

In this study, we investigate the crustal structure offshore north of Itu Aba Island (also known as Taiping Island) in the southern margin of the SW Sub-basin and the SW Sub-basin margin type by a combination of the MCS profile that intersects the published OBS profiles and the gravity modeling results. The MCS profile will reveal if the southern margin of the SW Sub-basin is more likely magma-poor or volcanic. The gravity modeling results will reveal the vertical variation of the crustal structure, providing a tentative re-examination of previous crustal structure interpretations of the

southern margin of the SW Sub-basin. The integration of the MCS and gravity modeling provides valuable insight into the crustal structure of the distal part of the southern SCS margin and the connections to regional tectonic evolution. Finally, to interpret the nature of the southern margin of the SCS, we construct a conceptual model of the HVDLC formation along with the SCS evolution, which includes the possible influence of a mantle plume and the hyper-extension of this wide and rifted margin.

2. Geological background The SCS oceanic basin is a marginal basin in the western Pacific with extremely wide rifted margins (Fig. 1). The origin of the SCS seafloor spreading has been explained by several models: (1) a back-arc basin probably caused by subduction of the Proto-SCS crust (Hsu, 1988), (2) a collisional basin related to crust extension due to the India-Eurasian collision and tectonic extrusion (Tapponnier et al., 1982), and (3) an oceanic basin related to continental rifting processes (Lewis, 2010). Recently, hybrid models have become more promising (Cullen, 2010; Morley, 2016). The seafloor spreading of the SCS oceanic basin was active during 32-15.5 Ma (Briais et al., 1993; Li et al., 2014; Chang et al., 2015a) or 32-20.5 Ma (Barckhausen et al., 2014; Morley, 2016). Although the origin and ages of the basin evolution are still questioned, it is widely accepted that continental breakup and seafloor spreading first occurred in the NW Sub-basin and were succeeded by at least one southward ridge jump and southwestward propagating rifting (Briais et al., 1993; Li et al., 2015; Qiu et al., 2016).

In addition to crustal evolution, the activities of the mantle plume beneath the crust are proposed to play a critical role in the opening of the SCS oceanic basin (Yan et al., 2007; 2014; 2015a; 2015b). In this hypothesis, the collision between India and Eurasia may initiate the nascent Hainan mantle plume during 50-30 Ma. In 30-16 Ma, the head of the nascent Hainan mantle plume had arrived in the asthenosphere and probably had interacted with the overlying lithosphere, enhancing the lithospheric deformation when the SCS oceanic basin was opening. After 16 Ma, the magmatic activities in the SCS oceanic basin were dominated by deep-sourced alkali basalt. Currently, a Hainan plume is imaged in the upper mantle, while originating in the lower mantle (Lei et al., 2009; Le et al., 2015; Zhao, 2015). It is characterized by a NW-SE tilting mantle low-velocity structure (Lei et al., 2009), with the tilt of the Hainan plume probably associated with the Manila subduction zone (Mériaux et al., 2015).

NW Sub-basin NW Sub-basin was formed as the earliest one among of the SCS Sub-basins. The seafloor spreading of the NW Sub-basin began around 32 Ma in a NW-SE direction, and ceased just after 30 Ma (Briais et al., 1993). It is bound by northern margin of the SCS to the north, and the Macclesfield Bank (Zhongsha Islands) to the south. Sedimentary successions in the NW Sub-basin consist of a syn-rift unit and a post-rift unit, overlying an oceanic crust of 5-7 km thick (Ding et al., 2011; Ding et al., 2012; Franke et al., 2014). The oceanic crust domain of the NW Sub-basin is characterized by highly reflective top of basement, little faulting and no discernible syn-tectonic strata, and the continental crust domain is characterized by tilted fault blocks overlain by syn-rift sediments and thins dramatically toward the oceanic crust domain

(Cameselle et al., 2015). Neither the SDR has been revealed by the MCS profiles nor the HVDLC has been found by OBS velocity profile in the NW Sub-basin and nearby margin area (Qiu et al., 2011; Ding et al., 2012), suggests that the margin of the NW Sub-basin is a magma-poor.

East Sub-basin East of the NW Sub-basin, the East Sub-basin occupies a vast area of the SCS oceanic basin. The seafloor spreading of the East Sub-basin lasted longer than in the other Sub-basins, starting from 32 Ma and ceasing around 15.5 Ma (Briais et al., 1993), with a jump event of the mid-ocean ridge (Qiu et al., 2016). It is bounded by the northern margin of the SCS to the north and by the southeast margin of the SCS off the Philippines to the south. The oceanic crust of the East Sub-basin is bounded and consumed by the Manila Trench to the east. The sedimentary successions consist of syn-spreading (Oligocene-middle Miocene) and post-spreading (late Miocene-present) deposits, based on IODP Expedition 349 drilling results (Li et al., 2015).

Whether the northern margin of the East Sub-basin is magma-poor, volcanic, or intermediate has been discussed for long time (Yan et al., 2001; Clift et al., 2001; Zhu et al., 2012; Gao et al., 2015). A HVL is generally identified in the lower crust along the northern margin of the East Sub-basin (Nissen et al., 1995; Yan et al., 2001; Wang et al., 2006), showing one of the diagnostic features of the volcanic margin. However, the MCS profile investigations suggest that although volcanic intrusions are widely recognized, the SDR, another diagnostic feature of the volcanic margin, have not been discovered in this area. Although igneous intrusions are extensively reported in the northern margin of the East Sub-basin (e.g., Sun et al., 2014), the absence of the SDR

reflection suggests that the northern margin of the East Sub-basin does not share similar characteristics with the volcanic margins in the northern North Atlantic (Mutter et al., 1982) and off northwest Australia (Hopper et al., 1992). An intermediate mode of margin type is therefore recognized (Zhu et al., 2012; Gao et al., 2015).

The volcanic complexes on the abyssal plain off the Philippines in the southeast margin of the East Sub-basin are distinct in bathymetric and seismic data (Franke et al., 2011). This volcanic province is confined to north of the Reed Bank north of NW Palawan and the Manila Trench. Based on the volcanic features revealed by Franke et al. (2011), the southern margin the East Sub-basin is more likely to be an intermediate margin. In addition, Franke et al. (2011) interpreted some low-angle detachments in the MCS profiles and suggested that hyperextended crust would occur in this area.

SW Sub-basin The SW Sub-basin is a V-shaped basin with the apex toward the southwest, revealing the southwestward propagation rifting. The seafloor spreading of the SW Sub-basin initiated at around 27 Ma and terminated at about 15.5 Ma (Briais et al., 1993), succeeded by the ridge jump of the spreading center. The earliest sedimentary succession of the SW Sub-basin is dated as middle Miocene, based on IODP Expedition 349 drilling results (Ding et al., 2016). The pre-Miocene sediments are mainly found along the base of the slope (Song and Li, 2015; Ding et al., 2013, 2016). The formation of the SW Sub-basin oceanic crust has been tentatively identified as hyperextended continental crust, exhumed subcontinental mantle, and steady-state oceanic crust (Savva et al., 2013; Ding et al., 2016).

The nature of the SW Sub-basin margin remains uncertain, especially on the southern margin. The northern margin of the SW Sub-basin is probably neither magma-poor nor volcanic (Lü et al., 2011). The earlier investigation of the southern margin of the SW Sub-basin (Line NH973-1, see Fig. 3 for the location) suggested that there was no HVL found in the area west of the Tizard Bank and Reefs (Zhenghe Reefs) (Qiu et al., 2011). Currently, a crustal structure with the "no HVDLC" result and a magma-poor margin are favored by some scholars (Chen et al., 2014). However, a later survey northeast of Line NH973-1 revealed several high velocity bodies beneath the lower crust (Line PR2, see Fig. 3 for the location) (Pichot et al., 2014).

3. Methodology A marine geology and geophysics (MG&G) survey was carried out by Taiwanese Research Vessel Ocean Researcher 1 in OR1-1068 cruise in 2014. The marine gravity data were collected by shipboard Micro-g LaCoste gravity meter Air-Sea System II. Marine seismic data were acquired by a 24-channel streamer, with 500 cube-inches air gun array shooting. Our seismic data (Line A in Fig. 3) intersect the published seismic profile of the Line NH973-1 (Ding et al., 2013; Song and Li, 2015) and OBS velocity profile of the Line PR2 (Pichot et al., 2014) (Fig. 5B). The seismic data were processed by ProMAX software, and the gravity modeling was performed by GeoModel program (see http://www.geoafrica.co.za/reddog/gc/geomodel/geomodel.htm for details).

To build up an intial model to perform the gravity modeling along the selected profiles, we separated the crustal structure into sediments, crust, and mantle. The

bathymetry and top of basement are derived from the MCS profile. We convert the time to depth with the water velocity (1500 m/s). The Moho depth along Line A is determined by the Parker-Oldenburg iterative method (Parker, 1973; Goḿez-Ortiza and Agarwal, 2005; Hsieh et al., 2010) with the global Bouguer anomaly dataset WGM2012 derived from International Gravimetric Bureau (see http://bgi.omp.obs-mip.fr/data-products/Grids-and-models/wgm2012 for details)(Bonvalot et al., 2012), and with regional average Moho depths referring to the Moho depth map by Braitenberg et al. (2006) and Li et al. (2010). The Moho depth of profile A in the Moho map of Braitenberg et al. (2006) is 13 km at both end with16 km at middle part, and 15-20 km in that of Li et al. (2010). We thus apply 15 km to the average Moho depth to perform the modeling. Density parameters of our layered crustal model used in this study are referred to Gao et al. (2015): the water (1.03 g/cm3), sediments (2.4 g/cm3), continental crust (2.7 g/cm3), HVDLC (2.97 g/cm3), and mantle (3.2 g/cm3).

4. Result In the MCS profile of the Line A, an acoustic boundary that separates the acoustic basement and the overlying sediments is cut by basement faults and features extrusive and intrusive structure highs (Fig. 4C). The basement faults are important features found in our seismic profile, forming several fault blocks as basement highs. Among these fault blocks, the northern extension of the Tizard Bank and Reefs (Zhenghe Reefs), which is located between the Western Taiping Seamount Group and the Zhenghe-Daoming Trough, is covered by high-amplitude, Oligocene-Miocene carbonate platform deposits. The nature of this carbonate platform deposit has been discussed in Chang et al. (2015b). Besides, there is an eastward updip reflection with

strong amplitude underlain by a group of concave-upward, antiform reflections in the central part of the profile (Fig. 5B). This eastward updip reflection may be also interpreted as a fault surface which is generated along with the overlying antiform structure.

The extrusive and intrusive structure highs are identified in the western and central-east part of the seismic profile, respectively. The presence of the extrusive structure is reflected in bathymetry, featuring two bathymetrical highs as the Western Taiping Seamount Group (Fig. 4 and 5A). The intrusive structures are recognized as buried highs as at the Zhenghe-Daoming Trough (Fig. 4, 5C and 5D). In the Western Taiping Seamount Group (Fig. 5A), The relief of the bathymetry is considerably parallel to the top of acoustic basement in the western part of the seismic profile, representing a veneer of sediments drastically uplifted by underlying volcanic basement structures. Besides, the intrusive structure high in Fig. 5D is bounded by the basement faults to its both east and west sides. This distribution pattern may reflect the development of basement faulting and volcanic intrusions.

Our seismic data also provide a re-examination to the southern part of the published seismic profile Line NH973-1 (Ding et al., 2013; Song and Li, 2015) and the OBS velocity profile Line PR2 (Pichot et al., 2014) (Fig. 3 and 4). The Line A intersects the Line NH973-1 at the Western Taiping Seamount Group (Fig. 4B and 5A), and intersects the Line PR2 at an intrusive structural high at the eastern part of our seismic profile (Fig. 4B and 5D). The previous investigations appeared to interpret the Western Taiping Seamount Group as fault block structures bounding a graben to the south (Ding et al., 2013; Song and Li, 2015). Based on seismic features of the

deformed sediments and basement structures in our seismic profile, we interpret the Western Taiping Seamount Group as the extrusive structures.

The gravity modeling along Line A revealed that there are HVDLC beneath the extrusive and intrusive structural highs observed in our MCS profile (Fig. 6). At the western part of the gravity modeling profile, a HVDLC is deeply seated about 10 km beneath the extrusive structure, with a 7-8 km in its thickest part. The thickness of this decreases eastward and this HVDLC eventually dies out in the central part of the profile. Further east, another HVDLC occurs beneath the intrusive basement structures in the central-east part of the profile. The thickest part of the eastern HVDLC is about 6-7 km, a little bit thinner than that of the western one. The eastern one extends farther east, out of range of our profile.

5. Discussion 5.1 Crustal structure north of Itu Aba Island (Taiping Island) With a lack of marine geology and geophysics (MG&G) survey data, the nature of the southern margin of the SW Sub-basin is less studied than that of the sub-basin margin. The crustal structure here remains unclear. A recent seismic investigation along the dip of the margin west of the Tizard Bank and Reefs (Zhenghe Reefs) by Qiu et al. (2011) presented a COT without HVL underneath the continental crust. A more recent seismic survey along the dip of the margin east of the Tizard Bank and Reefs by Pichot et al. (2014), however, suggested that some HVL bodies are imaged at the COT and continental crust. The difference in crustal structures occurs in such a short distance between both sides of the Tizard Bank and Reefs (Fig. 3) that a MCS profile

and gravity modeling results are prepared to better understand the crustal structure of the Tizard Bank and Reefs.

In addition to the fault blocks, our MCS profile documents two groups of volcanic basement structures, including extrusive structures in the Western Taiping Seamount Group in the western part of the profile (Fig. 5A) and intrusive structures in the Zhenghe-Daoming Trough in the central-eastern part of the profile (Fig. 5C). In contrast, gravity modeling results reveal the existence of the HVDLC in this area (Fig. 6). Furthermore, we found that the HVDLC occurs beneath the extrusive and intrusive structures, suggesting that the formations of these structures probably correspond to the presence of the HVDLC so that the extrusive and intrusive material may be injected from deep-seat HVDLC. We conclude that the southern margin of the SW Sub-basin is dominated by the volcanic basement structures and the HVDLC.

The SDR is generally considered as the diagnostic feature for volcanic margins. The absence of the SDR, along with the existence of the volcanic basement structures and underlying HVDLC shown in our data, suggest that the southern margin of the SW Sub-basin is probably similar to the margin of the East Sub-basin. Collectively, these features suggest that the southern margin of the SW Sub-basin is an intermediate type, similar to those in the northern margin of the SCS and in the margin north of NW Palawan.

5.2 HVDLC in the SCS 5.2.1. Magma-poor to intermediate transition between Sub-basin margin types

As noted above, the SCS oceanic basin is composed of the NW Sub-basin, East Sub-basin, and SW Sub-basin. These sub-basins were opened and developed serially: NW Sub-basin in 32-30 Ma, East Sub-basin in 32-16.5 Ma, and SW Sub-basin in 27-16.5 Ma (Briais et al. 1993). Also as noted above, the margins of the NW Sub-basin are shown to be magma-poor (Ding et al., 2012), and the margins of the East Sub-basin feature a HVDLC zone and extensive volcanic intrusive rocks (Gao et al., 2015), which is evidence of an intermediate type. This study revealed the margin of the SW Sub-basin to be intermediate, as are those in the East Sub-basin. This transitional feature has also been reported by Franke (2013) and Gao et al. (2016). In this case, an increase of magmatism is observed between the NW Sub-basin and the East Sub-basin. In addition, the intermediate mode observed in both margins of the East Sub-basin and the SW Sub-basin suggests that they probably bear a similar degree of magmatism. With this observation, it is suggested that both the HVDLC of the East and SW Sub-basin may have been formed together before the sub-basin started to spread or serially during the sub-basin opening.

Considering the timing of the sub-basin formation, it appears that the margin type of the sub-basins may change in response to the stepwise growth of the SCS oceanic basin. The early stage of the SCS oceanic basin evolution, as the opening of the NW and East Sub-basin began to form, is non-volcanic. The NW Sub-basin ceased to spread soon after 30 Ma, while the East Sub-basin kept spreading. At this time, a transition from the magma-poor margin to the intermediate margin of the East Sub-basin occurred, resulting in the HVDLC in the margin of the East Sub-basin. Simultaneously, the HVDLC in the SW Sub-basin and its margin formed as well.

Afterwards, the SW Sub-basin began to spread following a ridge jump, and the continued seafloor spreading led to the SCS oceanic basin in its present form.

A similar case of margin type change has been recognized off eastern Canada, where an along-strike change occurs from a volcanic margin in the south to a magma-poor margin in the north (Keen and Potter, 1995). This change has been ascribed to small-scale convection, which delivers large volumes of basaltic material. According to this concept, the margin transition from magma-poor to intermediate between the Sub-basins in the SCS, and the ridge jump related to the cessation of the NW Sub-basin and opening of SW Sub-basin are likely to be attributed to the addition of an active mantle-involved process.

5.2.2. Post-rift and post-spreading volcanism In the South China Sea, Cenozoic magmatisms are scattered, and the origin of some of the magmatisms remain unclear (Franke, 2013). In the Pratas Islands (Dongsha Islands) of the northern margin of the SCS, volcanic intrusions observed in reflection seismic profiles are proposed to postdate the Miocene sediments (Lüdmann and Wong, 1999). According to the seismic data, igneous bodies penetrated the Miocene or Upper Miocene deposits. To the west, a large volcanic zone along the distal northern SCS margin may represent a high extension zone and the intrusions may be attributed to lithospheric extension, seafloor spreading, or upwelling of the deep mantle material (Zhu et al., 2012). In the southern margin of the SCS, middle Miocene to Pleistocene sills, dykes, and plutonic materials are observed to occur near the Spratly Islands (Nansha Islands) area (Schlüter et al., 1996). Seafloor basalt samples collected along the toe-of-slope of the southern margin of the SCS are very young in age (~0.5 Ma)

(Kudrass et al., 1986). In addition, there were ~2 Ma volcanic ashes reported southwest of the Spratly Islands (IODP site 1143 in the Fig. 2; Shipboard Scientific Party, 2000). After summarizing the published age results in the SCS, Yan et al. (2006) suggested that the magmatic activities were very limited and most probably occurred after the cessation of seafloor spreading.

Prior to the post-spreading volcanisms, the post-rift volcanisms received increasing attention. Based on exploration wells and scientific drilling results, the ages of the volcanic events have been identified as 41.2, 27.2 and < 1Ma, suggesting that these events belonged to rift, post-rift, and post-spreading, respectively (Gao et al., 2015). Recent 40Ar/39Ar dating results from seafloor dredge samples in the northeastern SCS (SW of Taiwan) and from shallow drilling on Daimao Seamount in the East Sub-basin revealed the ages of 22-21 Ma and 16.6 Ma (Wang et al., 2012; Yan et al., 2015). A reflection seismic data analysis recently carried out by Zhao et al. (2016) shows that two groups of volcanic complexes developed during the late stage of the seafloor spreading. These post-rift volcanisms suggested that continental breakup will be a long-lasting event rather than just ceasing at particular time. To summarize, the post-rift volcanisms may be attributed to the breakup process, while the post-spreading volcanism is still enigmatic and requires more investigation and detailed analyses.

5.2.3. Nature and possible origin of the HVDLC in the SCS The HVDLC is one of the most important geophysical features in the margins of the SCS and may provide the most significant clues to the evolution of the margins and oceanic basin. The understanding of the crustal structures of the northern margin of

the SCS has been developed by numerous geophysical investigations, especially by the MCS and OBS investigations (McIntosh et al., 2014; Lester et al., 2014; Wang et al., 2006; Nissen et al., 1995; Yan et al., 2001; Wei et al., 2011; Gao et al., 2015). After reviewing these studies, Gao et al. (2015) summarized the extensive distribution of the HVDLC beneath the north margin of the SCS. To better present the distribution of the HVDLC in the northern margin of the SCS, a HVDLC thickness map is prepared based on these publications (Fig. 8).

In the thickness map, the HVDLC is widely distributed along the continental slope. The thickest part of the HVDLC in the northern margin of the SCS is evident in the Pratas Islands (Donsha Island) area, which is more than 12 km thick and generally thicker than 6 km. The thickness of the HVDLC tapers westward definitely, according to the MCS profiles provide by Gao et al. (2015) and the OBS profile provided by Yan et al. (2001). The eastward tapering is not that certain, since a local thick area marked by a contour line of 4 km based on the gridding result is likely to occur. The area of the HVDLC is estimated to be 1.2 x 105 km2. As an estimate, if we assume that the average thickness of the HVDLC is 5 km, the total volume of the HVDLC will reach more than 6 x 105 km3. It is much less than the thick igneous section off the east coast of the United States (3.2 x 106 km3 or 2.7 x 106 km3) (Holbrook and Kelemen, 1993; Kelemen and Holbrook, 1995). We also note that there is a high density block recently reported in the Taiwan-Luzon convergence belt, which is about 100 km east of the easternmost location in our mapping area (Doo et al., 2015).

Mantle serpentinization may account for the lower crust materials (Savva et al. 2013). However, to produce such a large extent for the HVDLC, underplating is believed to

be a more reasonable explanation than the serpentinization of the upper mantle material, which is generally limited to spatially occur along the distal margin (Pichot et al., 2014; Gao et al., 2015). In addition to serpentinization, the processes controlling the thick accumulations of higher velocity lower crust may fall into three distinct classes: (1) by rifting above abnormally hot mantle, (2) by active small-scale mantle convection of lower temperature mantle under the rift, or (3) by a fertile mantle (White et al., 2008).

On the basis of the early, limited ESP seismic data in the northern margin of the SCS, Nissen et al. (1995) considered that the high velocity layer may mainly result from inherited structures that are only partly related to the magmatic underplating. Based on the small extent of the high velocity layer discovered later and the possible coupling of the high velocity layer and the Plio/Pleistocene volcanism events, Yan et al. (2001) believed that the northern margin of the SCS is non-volcanic and that the limited high velocity layer is more likely related to hot, but not abnormally hot, mantle upwelling after the seafloor spreading event had ended. In addition, Yan et al. (2001) considered that the absence of the very rapid spreading rate, high heat flow, and extrusive volcanic sequence gave support for a non-volcanic origin margin. Recently, Gao et al. (2015) proposed that the underplating of the high velocity layer was in association with partially melting magma caused by the decompression of a passive, upwelling asthenosphere probably formed during early rifting or post-rifting. Similar process had been proposed by Savva et al. (2013). While a very limited amount of magma intruded upwards into the sedimentary basins, most magma had pooled at the base of the continental crust, forming the HVDLC in the northern margin of the East Sub-basin.

In addition to the passive mode, a mantle plume may provide explanations for the existence of the high velocity layer and a much more active basin opening style to explain the ridge jump. As Briais et al. (1993) mentioned, the ridge jump probably reflected the interactions between collisional extrusions of the India-Eurasian convergence and slab pull by the subducted Proto-SCS plate, and the change in the geodynamic boundary condition triggered both the jump and the propagation. Earlier, a linkage between the opening of the SCS and a mantle plume had been proposed by Deng et al. (1998). A later numerical simulation study suggested that additional lithospheric thinning related to the mantle plume is required to initiate seafloor spreading (Xia et al., 2006). However, if the mantle plume can play a major role in facilitating or driving continental rifting, the breakup of the SCS is often questioned (Flower, 1988; Yan et al., 2006).

The possible influence of a nascent Hainan mantle plume during seafloor spreading of the SCS oceanic basin (32-16 Ma) has received increasing attention (Wang, 2012; Sun et al., 2016). Geochemical studies of the young (3.8 - 7.9 Ma) alkali basalts in the SCS revealed that they were produced by a binary mixing between a depleted mid-ocean ridge basalt (MORB) mantle end-member and a type of enriched mantle end-member, indicating a plume origin (Yan et al., 2007, 2014, 2015). In addition, Yan et al. (2014) considered that a nascent Hainan mantle plume was likely to account for the ridge jump and, possibly, for the high velocity layer at a depth of 60-80 km under the SCS oceanic basin. This is probably the result of strong lateral mantle flows in association with the collisional extrusion of the India-Eurasian convergence and with slab pull related to the subducted proto-SCS plate. Similarly, Cullen et al. (2010)

introduced a splash plume model, in which a plume can result from instability caused by small pieces of downwelling slabs, and considered that it may account for some anomalous patterns in the SCS. This model appears to be able to provide some explanation of the nascent Hainan mantle plume, such as the short-life pulse crustal deformation from only 30-26 Ma. However, the direct evidence for the nascent Hainan mantle plume affecting the evolution of the SCS oceanic basin is still insufficient.

In addition to the passive mantle upwelling model and the plume model mentioned above, Lin et al. (2003) suggested that the higher velocity material in the northern margin of the SCS may be related to small-scale mantle convection, which is considered to evolve from instabilities of the cold and dense thermal boundary layer and to serve as a good candidate for providing the addition of heat flow (Ballmer et al., 2010; Sleep, 2011). They attributed the margin-scale uplift and vigorous magmatism during the rift-drift transition to the addition of the heat and its subsequent decay to the small-scale mantle convection. In this study, we suggested that such a large volume of the HVDLC should be closely related to the formation of the SCS oceanic basin and that the ridge jump should be interpreted by an even more active mode than the passive opening. Furthermore, more petrologic, geochemical, and geodynamic analyses to support a deep-seated, nascent Hainan mantle plume are still required. The small-scale mantle convection model, which is more active than the passive opening model and more passive than the mantle plume model, may provide an intermediate solution for the formation of the SCS oceanic basin.

5.2.4. Implication from volcanic basement structures and HVDLC

Compared to the earlier study in the Spratly Islands (Nansha Islands) south of our study area, which presented the seismic profile with non-volcanic interpretations (Hutchison and Vijiyan, 2010), our seismic profile in the northern Spratly Islands is dominated by both fault blocks and volcanic basement structures. Apparently, this distribution pattern of scattered fault blocks and volcanic basement structures is in association with the margin evolution. The fault blocks in our profile are mostly covered by the sediments. The volcanic basement structures are exposed mostly to the seafloor, indicating the volcanic basement structures are a lot younger in age (Fig. 5). In addition, the volcanic basement structures are observed to occur overlying the HVDLC (Fig. 6). Since fault blocks, volcanic basement structures, and the HVDLV are predominant in our profile, we postulate that the volcanic rock may come from the deep-seated HVDLC and form along the basement faults, which were able to provide a weak belt as a conduit for the magma to intrude. Large offset faults that may serve as good conduits for magma to intrude are generally formed close to the COT (Lavier and Manatschal, 2006; Franke et al., 2014).

The cause of the volcanic activities in this area is worthy of further discussion. The regional doming in the central and northern Spratly Islands is observed to be accompanied by faulting and volcanic emplacement (Zhou et al., 1995). This doming was earlier interpreted to be caused by a crustal bulge associated with crustal downwarping to the southeast. Recently, Chang et al. (2015b) proposed that the forebulge related to the Palawan-Borneo thrust wedge may be responsible for some regional uplift and possible volcanism. In this way, the effects of the fold-thrust belt may also play a role in the formation of the volcanism. Since more than one factor

may result in the volcanism in our study area, more geochemical and age constraints need to be developed in future studies.

We also found an eastward updip reflection underlying a group of antiform reflections in the central part of the profile (Fig. 5B). In the previous study, they were considered as an anticline in association with a drowning fault block (Chang et al., 2015b). However, a study off the mid-Norwegian coast has pointed out that the pre-breakup and post-breakup compressional structures may be formed in a hyper-extension environment (Lundin and Doré, 2011). They are observed and interpreted as the deformation related to a weaker part of the lithosphere caused by the hyper-extension processes. In this sense, the antiform reflections along with the underlying eastward updip reflection may also support the idea of the hyper-extension in the southern margin of the SCS.

Peron-Pinvidic et al. (2013) proposed a rifted margin model that separates a rifted margin into different structural domains, reflecting the phases of margin evolution. Succeeding stretching and thinning phases and a hyper-extension phase in which extremely thin continental crust and large-scale detachment faults occur will be followed by a magmatic phase, which may be the key phase for distinguishing magma-poor and volcanic margins. We consider that the HVDLC in our study area and related crustal weakening may be formed first in a stretching and thinning phase; while antiform structures and possible crust-scale faulting may be formed in a hyper-extension phase. Moreover, the absence of the SDR may suggest a subsequent amagmatic phase, indicating a magma-poor opening of the SW Sub-basin. The weakened continental crusts, which were influenced by the ponded HVDLCs and

hyper-extension processes, will tend to be deformed more easily by the sublithospheric activities, facilitating the latest volcanic events.

5.2.5. Lateral flow and geographic extent of the active mantle upwelling event As mentioned in Section 5.2.1, both HVDLCs of the East and SW Sub-basin may be formed together before the sub-basin started to spread or be formed serially during the sub-basin opening. The former condition suggests a larger extent of the topmost part of the ascending mantle, containing at least the unopened East and SW Sub-basin so that the unopened East and SW Sub-basin can be influenced at this time. The latter condition, in contrast, indicates that the extent of the HVDLC distribution may be controlled by the lateral flow of the topmost part of the ascending mantle that thickened the crust locally.

Lateral flow is a very significant behavior in a plume head. When a plume head encounters the base of the lithosphere, it is expected to flatten out and flow against this barrier. Sleep (1997, 2006) suggested that relief on the base of the lithosphere acts as an upside down drainage pattern with an enclosed catchment for this flow. In this way, the ascending mantle material will preferentially collect beneath regions where the lithosphere is locally thinning, such as at a thermally subsiding rifted margin. Subsequently, it begins to pond and to cool gradually, forming the lower crust material.

The melts caused by the nascent Hainan mantle plume occupied a very large area and cooled soon after they occurred, forming the HVDLC first and subsequently being separated by seafloor spreading. Conversely, when the lateral flow is more prominent

and after the ascending of the nascent Hainan mantle plume and the acceleration of the SW Sub-basin rifting, the melt generated in response will start to flow and accumulate toward the rifted continental margin of the unopened SW Sub-basin, forming the widespread HVDLC. These can be determined after the location and extent of the nascent Hainan mantle plume is further identified.

The geographic extent of the mantle ascending event may be directly associated with the location where the breakup event first occurred to determine the area that was influenced and the present form of the SCS oceanic sub-basins. Two candidates for the possible location of this active mantle ascending event can be discussed: (1) the thickest part of the HVL of the East Sub-basin around the Pratas Islands (Dongsha Islands) area, which is currently at the continental shelf-slope of the northern margin of the SCS (orange color area in Fig. 7); and (2) the region of the ridge jump during 30 - 26 Ma, which is currently at the seafloor of the East Sub-basin (purple circle area in Fig. 8).

In the former condition, if the mantle ascended first at the Pratas Islands area, the surface radius should be at least ~200 km if it has to reach and affect the region of the ridge jump. In the latter condition, the radius of the surface influence can be much less (~50 km). The less radius condition appears more consistent with the model with the ascending plume being related to splashed mantle material since this mantle upwelling may not be a large scale event. In this way, the thickest HVL in Pratas Islands area could be tentatively explained as another result of the lateral flow of the ascending mantle material. In addition, the thickness of the HVL may be even thicker than that in the Pratas Islands area before the crust was stretched. The thickness decreased

drastically and the HVL diminished after the later stretching and seafloor spreading occurred.

5.2.6. A comprehensive model The observations and considerations that may have been associated with the evolution of the northern southern margin of the SW Sub-basin are summarized and include (1) the transition between the magma-poor margin of the NW Sub-basin and the intermediate margin of the East and SW Sub-basins; (2) the cessation of the seafloor spreading in the NW Sub-basin was succeeded by a major ridge jump event, which very likely reflected an active mantle upwelling event; (3) the extensive distribution of the HVDLC beneath the well-investigated northern margin of the East Sub-basin and as explored in the SW Sub-basin in this study; and (4) the basement faulting and volcanism that may be related to the HVDLC and hyper-extension. Collectively, a conceptual model encompassing these observations is suggested, along with a map illustrating the locations of the HVDLC with time (Fig. 8) and the schematic diagrams of the selected profiles W and E showing the evolution of the HVDLC (Fig. 9).

The early stage of the SCS oceanic basin evolution (Fig. 9A), as the NW and East Sub-basin began to open, is non-volcanic. Subsequently, an active mantle upwelling, probably caused by a splashed nascent Hainan mantle plume, immediately generated melt that underplated the lower crust of the margin of the current East Sub-basin and formed the HVDLC as it cooled gradually (Fig. 9B). At the same time, the active mantle upwelling forced the spreading ridge to jump southward. The rifting of the southern margin of the SCS probably accelerated until the SW Sub-basin began to develop. The ponded melt may flow along the relief of the base of the lithosphere,

draining upside down within the enclosed catchment beneath the East and SW Sub-basin and mechanically weakening the overlying continental crust (Fig. 9B). The East and SW Sub-basins continued to spread, dividing the HVDLCs individually as their margins separated accordingly (Fig. 9C). Afterwards, the hyper-extension in the distal margins may also be responsible for crustal faulting, which may be formed along the weakened zone of the continental crust so that the continental crust became even weaker (Fig. 9D). The subsequent volcanic events, therefore, may tend to occur along the weak zone, facilitating the formation of the volcanic basement structures and resulting in the current state of the margins of the SCS Sub-basins (Fig. 9E).

Both active mantle upwelling event and hyper-extension would help improve our understanding of the SCS basin evolution. The presence of the active mantle upwelling event provides not only an alternative mechanism to explain the formation of the HVDLC, but also an active mechanism to help interpret the ridge jump of the SCS spreading center. Essentially, the hyper-extension is very likely to occur in this very wide rifted margin and potentially provides an explanation for later basement structures. We expect constraints on active mantle upwelling as well as hyper-extension processes in the SCS margins can be further provided to give a more complete vision to the evolution of the SCS.

6. Conclusion Based on the multi-channel seismic (MCS) and gravity data north offshore Itu Aba Island (Taiping Island) in the Spratly Islands (Nansha Islands), we revisited the crustal structures in the northern part of the southern margin of the SW Sub-basin of the South China Sea (SCS). The MCS data north offshore Itu Aba Island (Taiping Island)

suggest that the basement structural highs in the southwest margin of the SCS are dominated by fault blocks, volcanic extrusive and intrusive structures that probably formed along with basement faults. The gravity modeling results reveal that these volcanic basement structures were probably in association with the high velocity or density lower crust (HVDLC), suggesting that the southern margin of the SW Sub-basin features as an intermediate margin.

Based on the tectonic features in the SCS margins, including (1) the transition between magma-poor margins of the NW Sub-basin and intermediate margins of the East and SW Sub-basins; (2) the cessation of the seafloor spreading in the NW Sub-basin was succeeded by a major ridge jump event, which was very likely to reflect an active mantle upwelling event; (3) the extensive distribution of the HVDLC beneath in the well-investigated northern margin of the East Sub-basin and as explored in the SW Sub-basin in this study; (4) basement faulting and volcanism shown in this study that may be related to the HVDLC and hyper-extension, we proposed that the margin of the SCS basin was magma-poor at the time it occurred. An active mantle upwelling caused by small-scale mantle convection or the possible influence of the nascent Hainan mantle plume was formed subsequently, resulting in the HVDLC beneath the current margins of the East and SW Sub-basins, forcing the ridge jump, and opening the SW Sub-basin. The East and SW Sub-basins continued to spread, latitudinally dividing their HVDLCs as the margins separated accordingly. Afterwards, the hyper-extension in the distal margins may also be responsible for crust weakening zone that the subsequent volcanic events could occur within. Both active mantle upwelling event and hyper-extension would help improve our understanding of the SCS basin evolution.

Acknowledgements We are deeply indebted to Guest Editor Dei Eslava for considering our work. Stimulating review comments from Andrew Cullen and an anonymous reviewer are greatly appreciated. We are grateful to all the crew members of the R/V Ocean Researcher 1 who participated the MG&G investigation cruise OR1-1068, and to science party members of the National Taiwan University who helped with marine geophysical data collection. We wish to thank Dr. Jui-Lin Chang, Dr. Shyh-Chin Lan of the GeoResource Research Center, National Cheng Kung University, and Mr. Yuan-Wei Li, and Mr. Jian-Ming Chen of the Chinese Petroleum Company, Taiwan, for their encouragement. We thank the Bureau of Mine, Ministry of Economic Affair, Taiwan, for the financial support.

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FIGURE CAPTIONS Figure 1. Schematic diagrams showing crustal structures of (A) magma-poor margin (B) volcanic margin and (C) intermediate margin. EM: exhumed mantle; SDR: seaward dipping reflector; HVDLC: high velocity or density lower crust. V: extrusive and intrusive basement.

Figure 2. Regional map showing the South China Sea (SCS) oceanic basin. The study area is marked by the box in the southern SCS. The SCS oceanic basin is composed of the NW-Subbasin, East Subbasin, and SW Subbasin. The dash line indicates the continent-ocean boundary (COB)(Briais et al., 1993). The green and orange dots mark the location of the Itu Aba Island (Taiping Island) and the ODP site, respectively. The red lines and arrows indicate the margins conjugated. The white letters represents the margin type. MP: magma-poor; IT: intermediate.

Figure 3. Map of the study area showing the location of survey line A and other published profile referred to. Survey line A is marked with the green color. The thin dash line indicates the extent of the Western Taiping Seamount Group. The thick dash line indicates the continent-ocean boundary (Briais et al., 1993). The green and orange dots mark the location of the Itu Aba Island (Taiping Island) and the ODP site, respectively.

Figure 4. The shipboard gravity anomaly profile (A), uninterpreted MCS profile (B), and interpreted MCS profile with the drawing of the acoustic basement and basement faults along the Line A (C). The V.E.(vertical exaggeration) given by horizontal scale / vertical scale is 10 km / ~500 m, which is about 20.

Figure 5. Close-up views of selected sections from the MCS profile along Line A in Fig. 4B. The V.E. is given by 1 km / 0.4 second (~300 m), which is about 3.3.

Figure 6. Gravity model along the Line A. Density parameters of our layered crustal model used in this study are: water (1.03 g/cm3), sediments (2.4 g/cm3), continental crust (2.7 g/cm3), HVDLC (2.97 g/cm3), and mantle (3.2 g/cm3).

Figure 7. Map showing high velocity or density lower crust (HVDLC) thickness. This was constructed by previous published MCS profiles and OBS velocity profiles, including ESP-E (Nissen et al., 1995), OBS1993 (Yan et al., 2001), Wang et al. (2006), OBS2006-3 (Wei et al., 2011), MGL0908-3 (McIntosh et al., 2014) MGL0905-10 (McIntosh et al., 2014), MGL0905-20 (Lester et al., 2014), and L2 and L3 (Gao et al., 2015).

Figure 8. Map showing the postulated locations of the nascent Hainan mantle plume (purple filled circle). Colored stripes and thin lines are locations of the spreading center and the COB at different basin evolution stages (Briais et al., 1993). Note at the ridge jump event is indicated by the purple arrow, probably occurred during 30-26 Ma.

Figure 9. Conceptual model showing the evolutional stages of the SCS oceanic basin, based on the development of the sub-basins individually. (A) The early stage of the SCS oceanic basin evolution, as the NW and East Sub-basin began to open, is non-volcanic. (B) Subsequently, an active mantle upewelling, probably caused by a

splashed nascent Hainan mantle plume, immediately generated melt that underplated the lower crust of the margin of the current East Sub-basin and formed the HVDLC as it cooled gradually. At the same time, the active mantle upwelling forced the spreading ridge to jump southward. The rifting of the southern margin of the SCS probably thus accelerated, until the SW Sub-basin began to develop. The ponded melt may flow along the relief of the base of the lithosphere, draining upside- down within the enclosed catchment beneath the East and SW Sub-basin and mechanically weakening the overlying continental crust. (C) The East and SW Sub-basins continued to spread, dividing the HVDLCs individually as their margins separated accordingly. (D) Afterwards, the hyper-extension in the distal margins may also be responsible for crustal faulting, which may be formed along the weakened zone of the continental crust so that the continental crust became even weaker. (E) The subsequent volcanic events, therefore, may tend to occur along the weak zone, facilitating the formation of the volcanic basement structures and resulting in the current state of the margins of the SCS Sub-basins.

Graphical abstract

Highlights

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We found extrusive and intrusive structures north of Itu Aba Island (Taiping Island).

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The volcanic basement structures are overlying lower crust high density material.

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Margins of the SCS SW Sub-basin may be of intermediate mode.

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A SCS evolution model with active mantle upwelling and hyper-extension is proposed.