Crust and upper mantle structure and its tectonic implications in the South China Sea and adjacent regions

Journal of Asian Earth Sciences 62 (2013) 510–525

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Crust and upper mantle structure and its tectonic implications in the South China Sea and adjacent regions Qunshu Tang a,⇑, Chan Zheng b a b

Key Laboratory of Marginal Sea Geology, South China Sea Institute of Oceanology, Chinese Academy of Sciences, Guangzhou 510301, China College of Informatics, South China Agricultural University, Guangzhou 510642, China

a r t i c l e

i n f o

Article history: Received 5 April 2012 Received in revised form 16 August 2012 Accepted 27 October 2012 Available online 10 November 2012 Keywords: South China Sea Rayleigh waves tomography Shear wave velocity Lithosphere

a b s t r a c t We present a 3D S-velocity model for the crust and upper mantle of the South China Sea and the surrounding regions, constrained from the analysis of over 12,000 of fundamental Rayleigh wave dispersion curves between 10 s and 150 s periods. The lateral resolution was found to vary from 2° to 4° with the increasing period over the study region. A robust scheme of Debayle and Sambridge allowed us to conduct the tomographic inversion efficiently for massive datasets. Group velocity maps varying with period show lateral heterogeneities, well related to the geological and tectonic features in the study region. The 3D S-velocity model was constructed from the 1D structure inversion of the tomographic group velocity dispersion curves at each node. The obtained average crustal structure is similar to the PREM model, while the average mantle velocity is typically lower than the global average. The complicated 3D structures reveal three prominent features correlated with geological divisions: sea basin regions, island and arc regions, and continental regions. The derived crustal and lithospheric thicknesses range from 15 to >50 km and from 60 to >140 km, respectively, with the thinnest in the South China Sea, the thickest in eastern Tibet and the Yangtze Block, and the medium in the South China Fold Belt, Indochina, and island arc regions. Our results further confirm that (1) a Mesozoic subduction zone, which is interpreted as the tectonic weak zone during the Paleogene, exists along the South China margin; (2) the influence of the Indochina extrusion along the Red River Fault is limited for the South China Sea region; (3) there is a slab remnant of the proto-South China Sea beneath Borneo. New findings suggest that the Mesozoic subduction zone should be built into any evolution model for the region, as well as the other two major tectonic boundaries of the Red River Fault and proto-South China Sea subduction zone. Ó 2012 Elsevier Ltd. All rights reserved.

1. Introduction The South China Sea (SCS) is the largest marginal sea of the West Pacific and is located within the area of plate triple-junction that connects the Pacific, Eurasian, and Indo-Australian plates. The SCS is a relatively young sea basin, having opened since 32 Ma and stopped spreading at 16 Ma (Briais et al., 1993). The SCS is surrounded by very complex geological units and boundaries (e.g., continental blocks, island arcs, marginal seas, oceans, subduction zones, sutures, and fault systems) with perplexing tectonic activities and evolution histories. Therefore, a high-resolution three-dimensional (3D) model of crust and upper mantle is critical to improve our understanding of the structure and tectonics in this region. The study region of interest covers the main SE Asia (Fig. 1). Its geological framework experienced extensive modifications from the Paleozoic to Cenozoic (Hall, 1997; Metcalfe, 2009). By contrast, ⇑ Corresponding author. E-mail address: [email protected] (Q. Tang). 1367-9120/$ - see front matter Ó 2012 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.jseaes.2012.10.037

the Cenozoic evolution of SE Asia is the most concerned and extensively studied issue, especially in deciphering the opening of the SCS. Competing views about its origin have arisen for decades, and now fall into two main types: one is extrusion of the Indochina and the other is subduction of the proto-SCS (e.g., Tapponnier et al., 1986, 1990; Briais et al., 1993; Hall, 1997; Leloup et al., 2001; Morley, 2002; Searle, 2006; Clift et al., 2008; Hall et al., 2008). Irrespective of the evolution models considered, their related tectonic events dominating the regional tectonics must have left their imprints on the crust and upper mantle structures. In the present study, we seek to obtain the crust and upper mantle velocity structure, determine the geometry of the tectonic blocks, and provide seismic evidences on the lithospheric scale for unraveling the regional tectonic processes. To derive the 3D crust and upper mantle structure beneath SE Asia, we consider the regional surface wave tomography to be the optimal choice for the following two reasons: (1) body wave methods are strongly hampered by the sparse and uneven geographical distribution of seismic stations, which always fail to satisfy the density condition of crossing rays near the Earth

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Fig. 1. Simplified tectonic configuration of SE Asia. Abbreviations from the top of the diagram are: PRMB, Peal River Mouth Basin; YB, Yinggehai Basin; PI, Paracel Islands; MB, Macclesfield Bank; PKB, Phu Khanh Basin; RB, Reed Bank; and DG, Dangerous Grounds; PN, Panay Island; CR: Cagayan Ridge. The data of the tectonic boundaries are from the website: www.ig.utexas.edu. The 3000-m isobath of the South China Sea is contoured.

surface; (2) subduction zones surrounding SE Asia furnish sufficient earthquakes to ensure a good coverage of surface wave paths. Hence, SE Asia has always been a main- or sub-region in many previous works (Cao et al., 2001; Huang et al., 2003; Lebedev and Nolet, 2003; Okabe et al., 2004; Wu et al., 2004; Zhu, 2007). Cao et al. (2001) used the waveform inversion of 328 Rayleigh wave data from 4 broadband stations to study the crust and upper mantle shear velocity structures in the SCS and adjacent regions. Nearly in the same region, Wu et al. (2004) collected much more records of both Rayleigh and Love waves to acquire shear velocity structure. They also derived the distribution of crustal and lithospheric thicknesses from their 3D velocity model. Lebedev and Nolet (2003) studied upper mantle structure beneath SE Asia and West Pacific by the inversion of multimode surface waves, which have good constraints on deeper structures. In this paper, we have also presented a 3D S-velocity model of the crust and upper mantle in the SCS and adjacent regions from Rayleigh wave tomography. To achieve the result, a conventional two-step method has been used. First, geographical distribution maps of 2D group velocities at each period were inverted from the picked Rayleigh wave dispersion data. Then, group velocity dispersion curves at each node were inverted to construct a 3D velocity model. To better image the velocity structure of the region, we (1) selected more data than the previous works (Cao et al., 2001; Wu et al., 2004) to ensure a better path coverage; (2) introduced more records with short epicentral distances (10 ± 2°) that have narrower Fresnel zones to improve the lateral resolution; (3) applied an efficient tomographic algorithm (Debayle and Sambridge, 2004) that is suitable for extremely large datasets.

2. Data and method 2.1. Raw seismic data The vertical component seismographic data were used to determine the fundamental-mode Rayleigh wave group velocity dispersion curves. The data of the study region (latitude: 13°S–30°N, longitude: 90°E–130°E) were provided by 55 IRIS (Incorporated Research Institutions for Seismology, from January 1995 to February 2010) stations and 10 NDSN (National Digital Seismic Network, China, from January 2001 to December 2009) stations. Fig. 2 shows the locations of the seismic stations and over 12,000 earthquakes that occurred from 1995 to 2010. To ensure a better excitation of fundamental surface waves, only the earthquakes of magnitude greater than 5.0, epicentral distance greater than 5.0°, and source depth less than 100 km were retained. Stations with few seismic data were excluded, such as some temporary stations or newly installed stations.

2.2. Group velocity measurements After removing the instrumental responses, the vertical component seismograms of Rayleigh wave were analyzed to determine the group velocity dispersion curves between periods of 10 s and 150 s. The group velocity is determined following a multiple filter technique (MFT; Dziewonski et al., 1969) using an open source program, Computer Program in Seismology (CPS; Herrmann and Ammon, 2002). For extracting the pure fundamental-mode

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Rayleigh waves, the phase-matched filter (PMF; Herrin and Goforth, 1977) was also adopted to determine the last dispersion curves. Following the data processing as mentioned above, the group velocity data with only good dispersive features were selected. Fig. 3 shows an example of MFT processing for the Rayleigh wave with periods from 10 s to 150 s. Fig. 4a shows the path coverage at 20 s period with a total of 12,111 ray paths, displaying a good coverage in the study region. However, it can be noted in Fig. 4b that the ray density is not uniformly distributed. There is a sparse path coverage in Borneo and its adjacent regions compared to the northern part of the study region, but better than the peripheral regions, such as the Philippine Sea, southwestern Burma, and Indian Ocean near the Sunda Trench, where the ray coverage is more uneven (Fig. 4a). A histogram of the ray paths at selected periods is shown in Fig. 5a. The data for almost all the periods yield more than several thousand paths. The average group velocity (Fig. 5b) at each period

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derived from all the dispersion data, was used as the initial background model for 2D group velocity tomographic inversion (Section 2.3) and checkerboard resolution test (Section 2.4). Furthermore, it was used to obtain the laterally homogeneous S-velocity structure in the region (Section 4.1). Fig. 5c shows the path length distribution at 20 s period. Many short paths are included. 2.3. Building group velocity maps Before building a 3D structure, a key step is to map the 2D group velocity at each period through a tomographic technique. The tomographic algorithm of Debayle and Sambridge (2004) was used to invert for lateral variation in group velocity. This algorithm is based on the continuous formulation of the inverse problem and the least squares criterion. It incorporates some sophisticated geometrical algorithms that dramatically increase the computational efficiency and render possible the inversion of extremely large datasets. In the tomographic inversion, two important a priori parameters control the lateral smoothing of the model: a correlation length Lcorr acts as a spatial filter, and a standard deviation rM controls the amplitude of the model perturbations. In this study, the tomographic models were discretized into regular grids of 1°  1° with the average group velocities as the initial velocities for each period. To provide a better a priori parameter of the correlation length Lcorr, it was selected in a range of 40–600 km depending on the wavelength of each period approximately. Thus the resulted ‘‘influence zone’’ (2.64  Lcorr, a truncated Gaussian at 30 dB from its maximum) acts as a crude but effective way to account for finite-size sensitivity zones. (Debayle and Sambridge, 2004; Sieminski et al., 2004; Maggi et al., 2006). As similar with many previous works (e.g., Maggi et al., 2006), the a priori standard deviations rM of 0.07 km/s was used. In this study, we found that various Lcorr and rM used in the tomographic inversion would not lead to significant changes in velocity distribution. Fig. 6 is an example shows that the overall pattern of anomalies are unchanged by increasing Lcorr by a factor of two. 2.4. Reliability of the tomographic maps A number of factors can improve the lateral resolution of surface wave tomography, such as dense path coverage, relatively

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short propagation path, and proper inversion scheme accounting for finite frequency effects (Sieminski et al., 2004; Priestley and

Tilmann, 2009). Our study met these three above-mentioned conditions: the dense path coverage (Fig. 4), short propagation path (Fig. 5c), and optimized inversion scheme considering crude finite frequency effects (Sieminski et al., 2004). In addition, as no appropriate initial group velocity model exists for the inversion of individual periods, the regional average group velocities were used as the reference models, which were regarded as the preferred choices. In the following steps, two types of resolution tests were conducted to examine the effect of path coverage in the study region. A ‘‘quality criterion’’ scheme based on Voronoi diagram (Debayle and Sambridge, 2004) was used to assess the resolution. It was designed based on the geometrical distribution of the ray paths in each cell without accounting for the real data, models, and any other parameters of the inversion. The size of the Voronoi cells is used to give an indication of the length scale of the structures that can be resolved, while their shape provides information on the variation of azimuthal resolution. Here, the study region was initially discretized into a regular mesh of 1°  1° as the initial Voronoi diagram. The optimized Voronoi diagram for the dataset at 20 s period is shown in Fig. 7a. A noteworthy fact is that most of the Voronoi cells keep the starting shape in the main region except for the surroundings of the SCS. This indicates the lateral resolution to reach to 1°  1°. The result for the paths at 120 s period (Fig. 7b) also shows that most of the study region has the lateral resolution of 100 km without considering the Fresnel zone effect, although the number of paths is only about a quarter of the number of paths for 20 s period (Fig. 5a). However, the threshold of the criterion purely depending on the geometrical distribution was always too low to describe the actual lateral resolution. The lateral resolution assessed by the Voronoi diagrams can be degraded by the parameters that imposed the inversion, such as the correlation length Lcorr (Maggi et al., 2006). We also used the traditional checkerboard test (Zhao et al., 1992) to evaluate the achievable resolution although the method has intrinsic limitations (Leveque et al., 1993). The input models were set with 2°  2° and 4°  4° constant velocity cells, in which the perturbation of ±0.3 km/s were alternately superimposed on the average group velocities of the individual periods. The synthetic traveltimes were calculated using the same event-receiver combinations as the observations. Fig. 8 shows the recovered checkerboard models at 20, 40, 80, and 120 s periods. Fig. 8a and b shows that the geometry of the input models with 2°  2° cells for the periods around 20 or 40 s can be well retrieved except at

Q. Tang, C. Zheng / Journal of Asian Earth Sciences 62 (2013) 510–525

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the regions surrounding Borneo and the peripheral of the maps, where the paths coverage is relatively poor. Moreover, the prevalence of paths along certain directions causes some streaks on the maps. Fig. 8c and d shows the extremely faithfully recovered geometry and amplitude of the checkerboard models with 4°  4° cells over the entire region. We safely conclude that the actual lateral resolution could be reached to 2°  2° for the shorter periods and higher than 4°  4° for the longer periods in the study region.

3. Group velocity maps Group velocity maps at all the selected periods were obtained using the optimized tomographic scheme by Debayle and Sambridge (2004). Some representative results at 15, 20, 30, 50, 80, and 120 s periods are presented in Fig. 9. It is known that the sampling depth of Rayleigh wave increases with increasing period, and group velocity for a particular period is mainly sensitive to the average shear velocity integrated throughout the one-wavelength depth range (Mitra et al., 2006). Thus, the heterogeneities of the group velocity maps are related to the geological features and local tectonics with certain depth ranges. The group velocity maps at 15, 20, and 30 s periods (Fig. 9a–c) are primarily influenced by the upper crust, middle crust, and

lower crust, respectively, according to the sensitivity kernel derived from the regional average velocity model at 20 s period (Fig. 10b). Nevertheless, such correspondences are not exact because of the distinct types of the continental/oceanic crust in the study region. For example, the sea basin regions on these maps must be influenced by the uppermost mantle inevitably. The most distinctive feature in these three maps is the fast anomaly beneath the sea basin of the SCS, as well as the other sea basins, such as the Andaman Sea, Sulu Sea and Celebes Sea. Another distinctive feature is the good geographical correlation between the slow anomalies and the (continental) islands/arcs: the relatively low velocity anomalies strictly outline the boundaries between the islands/arcs and oceans. This feature can also be easily traced at periods longer than 30 s. Such findings are consistent with previous studies (e.g., Huang et al., 2003; Chang et al., 2007). However, the anomalies beneath the continental regions are variable. From the map at 15 s (Fig. 9a), the resulting group velocity in the mountain areas is generally higher than the sedimentary basin areas. In Fig. 9b (20 s), the manifestation for the continent is reversed rapidly into a slow anomaly feature. At 30 s period (Fig. 9c), a significant slow anomaly appears at the junction region of Tibet, Yangtze, and Indochina, with a notable velocity decreasing gradient towards northwest. The maps at longer periods (Fig. 9d–f) reveal typical upper mantle structures, where the variations of group velocity anomalies become smoother and weaker. Correlations between the

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surface geological features and the velocity anomalies distributions are degraded gradually. In Fig. 9d (50 s), the significant slow anomaly at the junction region of Tibet, Yangtze, and Indochina persists and grows pervasively on the northwest corner of the study region. By contrast, the strength of the fast anomaly beneath the SCS reaches the highest around 50 s period and then starts to decrease. Between these two primary anomalies, there is a relatively high velocity zone beneath the South China Fold Belt and its shelf regions. In Fig. 9e (80 s), the geographical correlation between the slow anomalies and the islands/arcs has been broken substantially. Meanwhile, the slow anomaly at the junction region of Tibet, Yangtze, and Indochina has mainly disappeared. In addition, a fast anomaly grows and covers most of the South China region, although its amplitude has become quite weak (Fig. 9e and f).

4. Shear wave velocity structure 4.1. Structure inversion Lateral variations of group velocity maps are found to be related to the geological structures and regional tectonics within certain depth ranges, but not a specific depth. A rather complicated map of group velocity could be generated with mixed information from different depths of the Earth. Thus the lateral comparison of the anomalies from the group velocity maps might be unreasonable and misleading without incorporation of solid evidences. The essential of this problem is the nonlinear relationship between the group velocity and the Earth’s structure. To find the 3D S-velocity structure of the study region, we first obtained 1D structures from the group velocity dispersion curves at each node of 1°  1° grid, and then combined them into 3D Svelocities. The linearized inversion program of surf96 (Herrmann and Ammon, 2002) was used following the damped least-squares

scheme. During the inversion, the 1D velocity model was parameterized into multiple layers with variable layer thickness (10 km and 15 km per layer above and below 100-km depth, respectively) in the depth range of 0–400 km. The S-velocity was inverted, and then the P-velocity is updated from S-velocity using the fixed Possion’s ratio of the initial model; the new density is computed from the new P-velocity using the empirical Nafe– Drake relation (Nafe and Drake, 1963; Herrmann and Ammon, 2002). Both differential smoothing and damping parameters were adopted to control the model roughness and stability of the inversion. In view of the dependency of the initial model during the inversion, we first made effort to derive a regional average velocity model as the initial model for the inversion of all nodes. Fig. 10a shows the inverted velocity model from the regional average dispersion curve using a constant velocity input model of 4.5 km/s. A tight fit between the observations and synthetic dispersion curve calculated from the inverted model is shown in the lower left of the figure. Simultaneously, Fig. 10b presents the model-derived partial derivatives of the group velocity at periods of 20, 50, and 150 s, providing schematic outlines of the vertical sensitivity. It can be found that the velocity structure above 200 km can be well constrained. Further, taking the inverted regional average velocity model as the a priori model of all nodes, we tried to obtain the 3D S-velocity model of the entire study region. After numerous tests, we chose a differential smoothing inversion with following parameters: a high damping value of 10 and iteration times of 30, because they produced tight fits between the observations and synthetic dispersion curves. Moreover, we found that the inversion was considerably stable because general features of the testing results remained the same, especially for the structures above 200 km. Fig. 11 shows the constant-depth cross-sections of the 3D S-velocity model. Figs. 12 and 13 illustrate the meridian and parallel cross-sections

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Fig. 10. (a) Inversion of the regional average group velocity dispersion curve using a half-space model with a velocity of 4.5 km/s. The black line is the inverted velocity model. The dashed line is the average model of the study region from the CUB model (University of Colorado Boulder, http://ciei.colorado.edu/~nshapiro/MODEL/). The gray line is the PREM model. The inset shows the fit of the average dispersion data (gray dots with error bars) with the synthetic dispersion curve (black line) calculated from the inverted velocity model. (b) Normalized group velocity partial derivatives for periods of 20 (black solid), 50 (gray solid), and 150 s (black dashed) derived from the inverted velocity model in (a).

in every 5°, with their location indicated at the bottom right of each panel, respectively. Fig. 14 shows the Moho and lithosphere–asthenosphere boundary (LAB) estimated from the 3D Svelocity model.

4.2. One-D and three-D S-velocity models Fig. 10a shows the average S-velocity inverted from the regional average dispersion data. An average crustal thickness of 30 km is

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Fig. 11. Shear wave velocity heterogeneities at different depths as shown in the lower-right corner of each panel. Thick lines on (a–c) represent the Red River Fault. The 3000m isobath (gray) of the South China Sea is contoured on (b).

shown in the model. Beneath this, there exists a mantle lid with an S-velocity of 4.4 km/s extending to 100-km depth. A low velocity zone (LVZ) exists beneath the mantle lid to the depth of 160 km with a sharp velocity drop to 4.25 ± 0.05 km/s. We found that the crustal structure is similar to the CUB model (Ritzwoller et al., 2002; Shapiro and Ritzwoller, 2002) and comparable with PREM model (Dziewonski and Anderson, 1981). Nevertheless, an overall low velocity feature of the mantle relative to PREM model is obviously observed. The appearance of this feature can also be found in the average model of the study region from the CUB model. A complicated 3D S-wave velocity structure is presented on the horizontal cross-sections in Fig. 11. Simultaneously, the vertical cross-sections (Figs. 12 and 13) illustrate the absolute S-velocity structures to help to further trace the velocity anomalies and provide visualized recognitions of the vertical/lateral structures, e.g., crustal/mantle/asthenospheric undulations and slab subductions. In general, the 3D S-velocity structures are found to have strong correspondences with the primary geological features, e.g., crustal/mantle features of the sea basins beneath the SCS, Sulu Sea, Celebes Sea, and Andaman Sea; velocity anomaly distributions beneath the island and arc regions; the continental velocity variations beneath the Indochina, South China, and Tibet blocks.

The following sub-sections give a brief account of the primary anomalies geographically (Sections 4.2.1–4.2.3). 4.2.1. Sea basin regions Four sea basins have been structurally unveiled in the study region: the SCS, Sulu Sea, Celebes Sea, and Andaman Sea, among which the SCS has been observed to be the most prominent. At 10-km depth (Fig. 11a), its high velocity structure does not correlate very well with the shape of the sea basin. This may be from the relatively sparse path coverage and the topography/bathymetry effect on the group velocity measurements at short periods less than 15 s (Pillet et al., 1999). The velocity value at 20-km depth (Fig. 11b) presents that it has been reached to the uppermost mantle. The shape of the fast anomaly at the northeast part outlines the sea basin faithfully, while the shape at the southwest part is distorted from the continent-ocean transition. We also found that structures beneath the thinned crustal areas are different: the Paracel Islands and the Dangerous Grounds are characterized with high-velocity anomalies but notably lower than the velocities beneath the Macclesfield Bank, the Phu Khanh Basin, and the Reed Bank. It appears that the middle or lower crust of the Macclesfield Bank, the Phu Khanh Basin, and the Reed Bank has been replaced by its mantle underneath. We observed a long and narrow zone

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Fig. 12. Seven shear wave velocity sections along the longitude direction in each 5° as indicated in the lower-right panel. The color scale is irregular in order to enhance the velocity contrast in the mantle. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

with slightly high velocity extending southwestward from the southwest sea basin and passing through the Sunda Shelf. This is very likely to have been caused by the intrusion of the mantle along a hidden fracture zone, probably southwest extension of the dead spreading ridge. Generally, the amplitude of the fast anomaly beneath the SCS starts to decrease with increasing depth (Fig. 11c). Instead of the high-velocity anomaly, a prominent low-velocity anomaly has appeared in the SCS region at 70-km depth and extending downward (Fig. 11d–f). We believe the asthenosphere has been uplifted to a depth shallower than 70 km. These findings are consistent with the observations by Cao et al. (2001), but different from the results by Wu et al. (2004), in which the high-velocity anomaly has been reported to still exist at depth greater than 85 km and the low-velocity anomaly of the asthenosphere is not significant. From the vertical cross-sections (Figs. 12 and 13) beneath the SCS, a very thin crust, a thin and high-velocity mantle lid, and an uplifted asthenosphere can be clearly recognized. The vertical patterns of the velocity structure in the Sulu, Celebes, and Andaman seas are generally similar to those of the SCS (Fig. 11a–e), although their lateral scales and velocity perturbations are relatively small. On comparison of these three small sea basins, the velocity anomaly of the crust beneath the Andaman Sea is weaker than that of the other two basins (Fig. 11a), which

might have been laterally smeared by the distinct low velocity beneath the Andaman Island. Beneath their deep part, such smallscale velocity structures (Fig. 11f) are apt to be disturbed or altered by deep tectonics, such as subductions or mantle flows. 4.2.2. Island and arc regions The island and arc regions include a very long and narrow volcanic arc belt and a large continental island of Borneo. The distribution of the volcanic arc belt is closely along with the seismic zones (Fig. 2b) from Ryukyu, Taiwan, Luzon, Philippines, Sulawesi, Maluku, Lesser Sunda, Java, Sumatra, Andaman Islands, to the western Burma onshore (Fig. 1). Particularly, the western Burma, which is different from the eastern Burma divided by the Sagaing Fault, is covered with the subduction-related volcanic arc and young basins (Mitchell, 1993; Morley, 2002; Searle et al., 2007). Above 30 km, these regions are found to be typically dominated with notable low velocities of less than 3.9 km/s (Fig. 11a–c). Meanwhile, from the vertical cross-sections (Figs. 12 and 13), the distinct thickened crust with low-velocity anomalies and downward curved contours can be traced. We believe that these are the typical nature of the hot and young volcanic arcs. However, at the depth greater than 70 km underneath most of these regions, the anomalies are found to be totally changed into high velocities (Fig. 11d and e). The vertical sections (Figs. 12 and 13) clearly show that the fast anomalies

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Fig. 14. (a) Crustal thickness. The region with warm (cold) color shows a deep (shallow) crustal depth. (b) Depth of the seismic lithosphere–asthenosphere boundary (LAB). The region with warm (cold) color means a shallow (deep) burial depth of the LAB. Note a reversed color scale is used relative to (a). The 3000-m isobath (gray) of the South China Sea is contoured. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

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have their specific inclines, which agree well with the subduction of the slabs. Therefore, in addition to the seismic zones along the belt, we consider that subduction of the slabs beneath the volcanic arc belt are responsible for the fast anomalies (e.g., Lei et al., 2009b). Nevertheless, at the triple junction of Indo-Australian, Pacific, and SE Asia plates, i.e., the region around Sulawesi, Maluku, and Lesser Sunda, there is no clear correlation between high-velocity features and the arc belts. Short-lived or inactive nature of the convergence zones and complex tectonics of the region (Harris, 2003) may break such correspondence. As one part of the SE Asia block, Borneo is a relatively stable area with little or no seismicity and active volcanoes today (Hall et al., 2008). However, above 30 km beneath Borneo, the lowvelocity feature is found to be similar to that of the volcanic arc belt, especially at northwest Borneo (Fig. 11a–c). At the depth of 70 km (Fig. 11d), a fast anomaly is observed to dominate the north Borneo, differing from the slow anomaly under the south Borneo. In vertical cross-sections (Figs. 11e and 12e), it can be traced that a fast anomaly of 500-km length dips mainly southward down to 100 km beneath Borneo. This fast anomaly has been reported previously, and been interpreted as the subducted slab of the proto-SCS (Rangin et al., 1999; Wu et al., 2004). 4.2.3. Continental regions Beneath the continental regions, including South China, Tibet, and Indochina blocks, there are prominent features with high and low S-velocity anomalies dominating the upper 10 km (Fig. 11a) and its underneath (Fig. 11b and c), respectively. This vertical structure pattern is similar to the group velocity distribution presented in Section 3. At 30-km depth (Fig. 11c), the pervasive low velocity of crustal feature beneath the continent indicates that the overall crustal thickness is greater than 30 km, especially close to Tibet. This can be seen intuitively from the vertical sections (Figs. 12 and 13). The low velocity feature at eastern Tibet and Yangtze Block extends to the depth greater than 70 km (Fig. 11d). It changes into fast anomaly at 100-km depth (Fig. 11e) and continues deeper than 160 km (Fig. 12). Most of the South China Fold Belt and Indochina regions at 70 km are revealed as fast anomalies, indicating the strong seismic lid of the continental lithosphere (Figs. 11d, 12, and 13). Most distinctly, a north-dipping fast anomaly greater than 4.5 km/s is traced from the northern SCS underneath to the coastal region of the South China Fold Belt (Figs. 12 and 13). In the view of previous works (e.g., Li and Li, 2007; Zhou et al., 2008), it is possible that the presumed Mesozoic subducted slab has been imaged successfully. The Red River Fault (Fig. 1) is one of the major strike-slip faults in the study region separating the South China block to the north and Indochina to the south. The velocity structures show that contrasts across the Red River Fault may be limited to the crust (Fig. 11b and c). Laterally, the contrasts across the fault zone are gradually weakened from the NW side near Tibet to the SE side near the coast, and finally disappeared completely at the Yinggehai Basin. These characteristics are the crucial constraints for the vertical penetration and lateral extension of the Red River Fault, as well as for its tectonic implications. 4.3. Regional lithosphere and asthenosphere Significant lateral variations of both Moho discontinuity and LAB are expected from the 3D S-velocity model. They also provide an intuitive view of the structural differences among the tectonic blocks. Fig. 14 shows the depth of Moho and LAB relative to the sea level, derived on the base of the S-velocity criteria as presented below.

4.3.1. Crustal undulation To find the crustal thickness variation in the study region, we used the depth of the 4.00 km/s S-velocity contour to denote the Moho approximately. The S-velocity 4.00 km/s is an optimal criterion which is determined by searching the best fit between the Moho depths derived from receiver functions (Bai et al., 2010; Chen et al., 2010; Wang et al., 2010; Kieling et al., 2011) and the S-velocity contours derived from the S-velocity model. Thus the depth of the 4.00 km/s contour closely corresponds to Moho depth over much of the region. The overall variation of the Moho is similar to Mooney et al. (1998) and Wu et al. (2004). In the basin region of SCS, the burial depth of Moho (Fig. 14a) typically ranges from 15 km at the center to 20 km around the margins. Generally, these estimates are consistent with the results from gravity or seismic refraction data (e.g., Nissen et al., 1995; Xia, 1997; Li et al., 2010). It is noteworthy that the estimated Moho in the sea basin region would change slightly (1 km) if a water layer was considered. Along the coastal region of the South China and Indochina, there is a high gradient transition zone with 20–30 km Moho depth. To the north and west of this zone, the typical continental Moho with the burial depth ranging from 30 to 40 km is observed, matching exactly with the estimates by Wu et al. (2004) and other results (Zeng et al., 1997; Hu et al., 2005; Zhang and Wang, 2007; Chen et al., 2010). The crustal thickness near the eastern Tibet and Yangtze Block increases dramatically to greater than 50 km, which is very similar to the previous findings (e.g., Zhang et al., 2007; Yao et al., 2008). The crustal thickness is 30 km beneath the island and arc regions, notably deeper than their abutting sea areas. Unlike the thick crust beneath the southern Sunda Shelf, there is a distinct Moho uplift below the northern Sunda Shelf and southern Indochina. This feature obviously contradicts with the widespread assumption that the Sunda Shelf is a uniform and long-term stable region (e.g., Hutchison, 2004), which has been doubted previously from several individual aspects (Hall, 2002; Hall and Morley, 2004). Moreover, it seems that the Moho uplift connects with the SCS in the east and Andaman Sea in the west, and can be further traced along the Sagaing Fault, and north to the border between Burma and India. If this extrapolation is feasible, then the Moho uplift could be an important implication for an unknown terrane boundary beneath the Sunda Shelf based on the analogous constraints for the Mesozoic tectonic boundaries along the Sagaing Fault (Morley, 2002; Metcalfe, 2009) and northeastern SCS (Li et al., 2007; Zhou et al., 2008; Tang and Zheng, 2010). 4.3.2. Seismic LAB There are several ways to define the seismic LAB (e.g., Li and Burke, 2006). In the previous study by Wu et al. (2004), they took the maximum S-velocity as the bottom of the lithosphere. We consider that their definition is unreasonable and their results are not robust and representative. Another definition using the fixed S-velocity above the LVZ can also create uncertainties, not only because the choice of the absolute velocity is somewhat subjective, but also the velocities from different regions and depths would lose their comparability (An and Shi, 2006b). In the present study, we have employed the classical definition of the seismic LAB using the strongest negative vertical gradient above the LVZ (e.g., Priestley and Debayle, 2003; Artemieva, 2009). Fig. 14b shows the undulation of the LAB in the study region. It has been smoothed using a median filter to suppress the sharp variations. The lithosphere beneath the SCS region is distinctly thin varying from 60 to 80 km. Remarkably, there are two primary zones of LAB uplift up to 60-km depth at the northern and southern part of the sea region with W–E and NE–SW strike, respectively. The multi-phase spreading of the SCS may be responsible for the spatial distribution of the raised LAB zones. The northern LAB uplift

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extends from Luzon Strait to Xisha Trough. It is displaced from the northern spreading center of the SCS (Briais et al., 1993), but coincides well with the Mesozoic subduction or suture zone (e.g., Li et al., 2007). Uplifts in south are more uneven and segmented. The major uplift is about 200 km SE from the southern spreading center. Another smaller uplift is located at the southeast of the Vietnam margin. Similar to the northern uplift, southern uplifts are also offset from the spreading center. This may be resulted from complicated evolution process as discussed by Bastow et al. (2005). On the contrary, the lithosphere beneath the Sulu, Celebes, and Andaman Sea is observed to be in the range of normal thickness (100 km). Beneath the South China Fold Belt and Indochina, the lithosphere thickness is typically in the range of 80–100 km. Such findings are similar to those observed by Cao et al. (2001) and Huang et al. (2003), as well as the estimate of the seismic-thermal lithosphere (An and Shi, 2006a). The lithosphere thickens sharply beneath the Yangtze Block and eastern Tibet from 100 km to more than 140 km. Previous studies have revealed a much thicker lithosphere down to 170–200 km in this region (Huang et al., 2003; Zhu, 2007). In fact, the lithosphere thickness of our estimates has reached 180 km as presented in Fig. 12a and b, but these ‘‘singular values’’ have been smoothed by the spatial filter. Being similar to the Moho undulation, the LAB is also downcurved under the island and arc regions. The subducted slabs tend to affect the thickness estimation in these regions, and may produce a thicker lithosphere. By contrast, the LAB is uplifted to lower than 80 km along the east of the Sagaing Fault, Gulf of Thailand, and northeast of the Sumatra. It seems that there is a good spatial correlation between the thinned lithosphere and the subduction zone. One possibility is that the subduction-related lithospheric erosion in the mantle wedge causes thinned lithosphere (e.g., Macpherson, 2008). However, it is considered to be unreasonable because the uplift is not centered beneath the volcanic arcs. Another possibility is that the ‘‘thinned lithosphere’’ is regionally normal with a thickness of 80 km, and it is actually an illusion relative to the overestimated thickness along the subducted slabs as mentioned earlier. Alternatively, we can reconnect the uplifted zone from the east of the Sagaing Fault, Gulf of Thailand, Sunda Shelf, and to the west and northeast of SCS. This zone generally corresponds well to the uplifted Moho described in the last subsection, but to the 200–400 km horizontal offset along the west segment. As prompted by the pre-Cenozoic tectonics and many previous findings of this region (Liu et al., 2006; Li et al., 2007; Zhou et al., 2008; Metcalfe, 2011), we presume that a great preCenozoic suture might have existed along these regions, and became further curved or reformed during Cenozoic by northward pushing of the Indian plate, southward extrusion of the Indochina block, and southward extension of the SCS. 4.3.3. Seismic LVZ The seismic LVZ is ubiquitous beneath the lithosphere with reduced S-velocities of 4.2–4.3 km/s (Figs. 12 and 13). Similar to the large top undulation of the LVZ, the bottom fluctuation of the LVZ is also significant, making a wide thickness variation of the LVZ from 50 to 150 km. The most prominent LVZ is right beneath the SCS region. It is a lens-shaped LVZ extending down to 200km depth. No distinct anomaly could be recognized as the plume-like low velocity channel. This finding does not support the view of plume-related mantle upwelling, which is a proposed mechanism of the opening of the SCS (e.g., Wang et al., 1995; Li and Yang, 1997). It is noteworthy that evidence from our model does not support the existence of the well-known Hainan plume (80 km) mostly examined using body wave tomography (e.g., Montelli et al., 2004; Zhao, 2007; Lei et al., 2009a). This might be because our model does not have the lateral resolvability for the

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anomaly smaller than 80 km. There is an exception under the eastern Tibet and Yangtze Block, where the top of the LVZ can still be observed with weak vertical gradients around 180-km depth, but its bottom has been out of the scope (Fig. 12) and may extent down to a deeper region, as presented by Huang et al. (2003). 5. Tectonic evolution of SCS The high-resolution 3D S-velocity model unveils regional- to local-scale lithosphere/asthenosphere structures, and provides critical seismic evidences for understanding the tectonics of the study region. One of the most concerned issues is the tectonic evolution of the SCS, which has been extensively studied (e.g., Tapponnier et al., 1986; Hall, 1997; Leloup et al., 2001; Morley, 2002; Sun et al., 2006; Clift et al., 2008; Cullen et al., 2010). Two passive-rift models of the SCS opening have been proposed as the end-member models differing in the relative importance of extrusion versus subduction as the driving mechanism (Clift et al., 2008; Cullen et al., 2010). In view of the inconsistencies between the observations and the end-member models, hybrid models have been proposed to accommodate the conflicting views (Morley, 2002; Cullen et al., 2010). In the present study, based on the numerous previous findings and our new evidences, we have focused on several critical aspects strongly related to the opening of the SCS, and have put forward a new perspective of the regional tectonics. 5.1. The Mesozoic subduction zone There is a Mesozoic subduction or suture zone along the north margin of the SCS as reported in many previous works (e.g., Nissen et al., 1995; Sibuet et al., 2002; Liu et al., 2006; Li et al., 2007; Li and Li, 2007; Zhou et al., 2008; Tang and Zheng, 2010). Along this zone, a tectonic weak zone was easy be developed by the successive compression-uplifting, lithosphere shortening, and extension. However, all these findings and assumptions are limited by their evidences from the crustal structures and/or the local or distal geochemical samples. Our high-resolution velocity model clearly presents a fast anomaly extending from the north margin of the SCS beneath to the South China Fold Belt. A preferential interpretation of the fast anomaly is the remained Mesozoic flat-subducted slab. Such interpretation is extremely compatible with the previous views about the Mesozoic subduction zone. Therefore, when the tectonic regime changed from active to passive in the South China continental margin during the Paleogene (e.g., Taylor and Hayes, 1983; Li et al., 2007), this subduction zone would have acted as a tectonic weak zone that could be easily reactivated in a rifting system. Our observation of the thinned lithosphere or uplifted asthenosphere along the northern margin of the SCS (Fig. 14b) perfectly matches with the tectonic weak zone as expected. It has been assumed that once the lithospheric extension (e.g., Hayes et al., 1995; Rangin et al., 1995; Zhou et al., 1995) could not sustain the block displacement, a breakup along the weak zone would be developed. Thus, the South China margin has become a stress-free boundary. If this assumption is true, the successive extension until 25 Ma at the continental margin (Clift and Lin, 2001) should not be simply related to the slab-pull effect from the south, but the extrusion regime from the northwest. 5.2. The Red River Fault The Red River Fault is a reactivated Mesozoic fault zone in response to the India–Asia collision with more than 500 km of strike-slip offset during 32–17 Ma, which is synchronous with the seafloor spreading of the SCS (e.g., Leloup et al., 2001). Thus, the extrusion of the Indochina block relative to the static South

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China block along the Red River Fault is responsible for the SCS opening. However, our model shows that the Red River Fault is not a very prominent zone especially below the crust. The lateral variation of the velocity anomalies reveals a gradually weakened influence southeastward along the fault zone. Recent results from both surface wave analysis and shear-wave splitting have shown that the fast azimuthal anisotropy orientation turns from SE at northern Indochina to NE at the east Vietnam margin (Hao et al., 2005; Bai et al., 2009), strongly implying the stress field at the east Vietnam margin is no longer a sinistral strike-slip dominating feature (e.g., Roques et al., 1997). These observations indicate that the influence of the extrusion in SE Asia is not so tremendous as presented previously, as well as the offset of the Red River Fault (Rangin et al., 1995; Roques et al., 1997; Leloup et al., 2001; Morley, 2007; Clift et al., 2008; Hall et al., 2008). Nevertheless, the extrusion had been a dominant factor north of the east Vietnam margin until 25 Ma. The major extension of the South China continental margin as well as the first-stage seafloor spreading ceased at 25 Ma (Briais et al., 1993; Clift and Lin, 2001). Both continental extension and seafloor spreading could have been transferred to the left-lateral motion of the Red River Fault (e.g., Rangin et al., 1995; Leloup et al., 2001; Morley, 2002). Thus, the total offset of the Red River Fault near the SE extremity could not exceed 500 km: 300-km continental extension of the South China margin (Hayes et al., 1995) and 200-km seafloor spreading of the northern SCS basin. This amount is less than the estimates of 500–1000 km (Tapponnier et al., 1990; Leloup et al., 2001) and higher than the estimate of 280 km by Clift et al. (2008). Obviously, a 500-km displacement would be insufficient to trigger the southward ridge jump of the SCS as well as to further sustain the spreading of the southern sea basin, which is wider than the northern one. 5.3. The proto-SCS subduction It has been widely documented that there was an oceanic basin named proto-SCS between South China margin and Borneo (e.g., Rangin and Silver, 1991; Hall, 2002; Clift et al., 2008). The protoSCS subducted mainly southward beneath Borneo from 45 to 16 Ma (Clift et al., 2008; Hall et al., 2008). This is regarded as an alternative mechanism of regional tectonics that the slab-pull of the proto-SCS might create an extensional regime, and then cause the SCS opening. This geology-based model has been supported by strong geophysical evidences both from mantle and crustal structures (e.g., Rangin et al., 1999; Clift et al., 2008). Coincidentally, our result also reveals a 500-km fast anomaly dipping south-southeastward below northern Borneo (Figs. 12 and 13), longer than the estimate by Rangin et al. (1999). The preferential interpretation of the fast anomaly is the slab remnant of protoSCS. Therefore, we also believe that the slab-pull is an important factor for the evolution of the SCS. Considering the existence of the tectonic weak zone along the northern margin of the SCS, the influence zone of slab-pull would have been confined to its south region, causing the crustal thinning of the Dangerous Grounds and Reed Bank from Eocene to Oligocene (Cullen et al., 2010), and the southward movement of these blocks. Thus, at least two tectonic forces of extrusion and slab-pull would have coexisted during the first stage of seafloor spreading, and the former disappeared after 25 Ma. This may be the reason that for the spreading rate in the north basin is faster than that in the south one (Briais et al., 1993). Meanwhile, the seafloor spreading in the first phase might help to promote the subduction of the proto-SCS mutually. Geographically, the general strike of the new spreading center was changed to NE–SW, similar to that of the major LAB uplift observed to the southeast of the spreading center. Their strikes are

also in accordance with the subduction zone of the proto-SCS. These indicate that the slab-pull might have dominated the regional tectonics in the south. Therefore, we suppose that the southward ridge jump at 25 Ma (Wang and Li, 2009) and its successive spreading might be initiated and sustained by one stage of the rapid subduction during that period, which may coincide with the one phase of the episodic volcanic eruption (23– 15 Ma) as recorded on the Panay Island northeast of the Sulu Sea and the Cagayan Ridge in the Sulu Sea (Bellon and Rangin, 1991; Rangin and Silver, 1991). 5.4. A revised evolution model Our high-resolution 3D S-velocity model rendered critical seismic evidences for three main tectonic boundaries around the SCS: a buried Mesozoic subduction zone to the north, a giant strike-slip fault zone to the northwest, and a ceased Cenozoic subduction zone to the south. These findings enhance previous understanding of the regional evolution models which were considered to be controlled by these tectonic boundaries. Meanwhile, the role of these findings from the Earth’s deep structure needs to be reevaluated. Current models concerning the SCS evolution primary focus on extrusion of the Indochina and/or slab-pull of the proto-SCS, while the Mesozoic tectonic weak zone along the South China margin is seldom considered. Therefore, based on previous findings and our new observations, a revised evolution model of the SCS has been proposed (Fig. 15). The Paleogene extensional regime of the South China margin was simultaneously dominated by both India–Asia collision (primary) and proto-SCS subduction (secondary), causing the initial lithospheric extension. Because of the existence of the Mesozoic subduction zone, the successive extension caused the breakup of the nascent SCS along the weak zone at 32 Ma (Fig. 15a–c). The first-stage spreading before 25 Ma was sustained by the gradually weakened extrusion and strengthened slab-pull. As estimated, a maximum of 500-km extension north of the east Vietnam margin was transferred to the Red River Fault. Therefore, the role of the Mesozoic subduction zone is extremely significant in our reconstruction because it may control the strike of the northern spreading center and prevent the southeastward extension of the Red River Fault into the SCS. The south of northern spreading center, the tectonics was mainly controlled by the southeastward subduction of the protoSCS, resulting in the thinned crust in the Dangerous Grounds and Reed Bank regions. Meanwhile, the subduction was also aggravated by the southward propagation of the seafloor spreading. A rapid subduction during 25–16 Ma initiated the major jump of the spreading center and maintained the rifting in the southern SCS (Fig. 15d). The spreading was ceased until the continental microblocks of Dangerous Grounds and Reed Bank collided with Borneo along the North Borneo Trough (Fig. 15e). 6. Summary (1) Tomographic group velocity maps between periods 10 and 150 s were obtained from the analysis of more than 12,000 fundamental Rayleigh wave seismograms in the SCS and adjacent regions. The lateral resolution was estimated to be 2–4° throughout the study region. Group velocity maps showed systematic variations with increasing periods and good lateral consistencies with the known geological and tectonic features. (2) A 3D S-velocity model of the crust and upper mantle was constructed from the 1D inversion of the tomographic group velocity dispersion curves with a regional average velocity

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(a)

(b)

(d)

(e)

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(c)

Fig. 15. Tectonic reconstructions of the South China Sea and surrounding regions at 150 Ma, 40 Ma, 30 Ma, 20 Ma and 15 Ma. The base map of the model is derived from Hall (2002) and Li et al. (2007).

structure as the initial model. The crustal structure of the regional average velocity model is similar to the reference models, e.g., PREM model, while the mantle velocity is typically lower than the global average. The complicated 3D structures primarily fall into three prominent features in response to their geological divisions: the sea basin regions, the island and arc regions, and the continental regions. (3) Both Moho discontinuity and seismic lithosphere–asthenosphere boundary were derived as the indicators of the crust and lithosphere thicknesses. Their strong lateral undulations also correspond well to the geological divisions. The crust and lithosphere thicknesses ranged from 15 to >50 km and from 60 to >140 km, respectively, with the thinnest in SCS, the thickest in eastern Tibet and Yangtze Block, and the medial in South China Fold Belt, Indochina, and island/arc regions. The thickness of the seismic low velocity zone beneath the lithosphere was found to vary from 50 to 150 km. (4) A Mesozoic subduction zone has been interpreted as the tectonic weak zone during Paleogene along the South China margin. The influence of the Indochina extrusion along the Red River Fault is not yet as tremendous as presented previously. The view of proto-SCS subduction is enhanced by the new evidence of the slab remnant beneath Borneo. It is the dominating factor of the extensional regime from the Eocene to early Miocene. (5) A revised evolution model of the SCS is proposed. Being relative to the previous models, which are mainly concerned about the Red River Fault and proto-SCS, the tectonic weak zone along the South China margin has been considered. This provides insight into the initial opening of the SCS: spatial location (along the weak zone) and temporal constraint (first opening stage). Therefore, it is strongly recommended to build the Mesozoic subduction zone into the evolution model for the region.

Acknowledgments The IRIS (www.iris.edu) and NDSN (www.csndmc.ac.cn) seismic data management centers are highly appreciated. Two anonymous reviewers are greatly acknowledged for their constructive comments. We thank E. Debayle and M. Sambridge for their tomographic code (Debayle and Sambridge, 2004) and Professor Robert Herrmann for his software package of Computer Programs in Seismology (CPS; Herrmann and Ammon, 2002). Generic Mapping Tools (GMT; Wessel and Smith, 1998) is used for preparing all the figures. This research is supported financially by the National Natural Science Foundation of China (Grants 40904011, 41176026, and 40806022) and the Knowledge Innovation Program of the Chinese Academy of Sciences (Grant SQ200908).

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