Geological setting and tectonic evolution of the South China Sea

South China Sea is a marginal sea of Eurasia and borders the South China Block, Indochina Block, Luzon arc, Palawan. The opening of South China Sea is considered to have occurred during the Oligocene and is one of the youngest ocean basins in the world (Ludwig, 1970; Ben-Avraham and Uyeda, 1973; Briais et al., 1993; Lee and Lawver, 1995,). Based on the magnetic anomaly mapping in 1979 (anomalies 11 to 5d, Taylor and Hayes, 1980, 1983; revised to anomalies 11 to 5c, Briais et al., 1993), the opening lasted from the late Oligocene to the early Middle Miocene (32 to 15 Ma). The tectonic process responsible for the opening of the South China Sea is debated.

There are three principal models which describe the formation of the South China Sea. 1) The South China Sea is an inactive marginal basin and opened as a consequence of back-arc extension due to eastward subduction of the Eurasia plate (Karig, 1971; Ben-Avraham and Uyeda, 1973, Zhang, 1984). 2) The South China Sea is opened mainly due to the mantle lateral flow or mantle plume (Watanabe et al., 1977; Taylor and Hayes, 1980; Flower et al., 1998; Zhang et al., 2001). 3) The South China Sea opened as a consequence of regional displacement in SE Asia due to the collision of India-Eurasia at ~50 Ma (Tapponnier et al., 1982, 1986, 1990; Lee and Lawver, 1995).

Karig (1971) proposed that the South China Sea is perhaps an inactive marginal basin with higher than normal heat flow value than traditional marginal basin. Ben-Avraham and Uyeda (1973) mentioned the opening of the South China

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Sea may be related to the subduction of the Pacific plate and the compression of the China Sea basin. Zhang (1984) proposed that South China Sea as well as East China Sea are the continental margin under tension and the subduction of the plate led to the consequence of post-arc pull-apart basin. This interpretation was challenged when Taylor interpreted the magnetic anomalies. The magnetic anomalies reveal the E-W direction expansion axis of the South China Sea, which is unreasonable to be a backarc basin of the Luzon arc with the nearly N-S direction subduction axis (Taylor and Hayes, 1980).

Watanabe et al. (1977) published the heat flow of the South China Sea with the value 85 ± 8 mW/m² at north and 112 ± 16 mW/m² at south of the basin. In this paper (Watanabe et al., 1977), they also predicted an age of 14-36 Ma (lower Miocene to lower Oligocene) for the South China Basin. Taylor and Hayes (1980, 1983) proposed the magnetic anomaly of the spreading axis of South China Sea, and interpreted the spreading of the South China Sea between ~32 Ma and ~17 Ma (mid Oligocene through early Miocene). Flower et al. (1998) proposed a model explaining the dispersed volcanism and DUPAL-like asthenosphere in East Asia and West Pacific. Zhang et al. (2001) discussed the mantle lateral flow and the tectonic regime relationship and suggested that the lateral flow was the dominant factor of the thinning of the lithosphere and the expansion of the South China Sea. However, Cui et al. (2005) and Xia (2005) remodeled the convection of the mantle flow and suggested that the upwelling of the hot mantle material had an impact on the thinning of the lower lithosphere but less influence on the crust thinning and thus taking much longer time for the opening.

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Tapponnier et al. (1982, 1986, 1990) proposed that the collision of the India block and Eurasia had a large impact on the East China, South China and Indochina blocks. The collision pushed out the East China and South China block and dominated the left-lateral shearing of the Ailao Shan-Red River metamorphic belt.

They proposed that the penetration of India block into the Eurasian continent pushed Indochina southeastward with clockwise rotation, and the over 500 km left-lateral movement directly influenced the Oligocene-Miocene opening of the South China Sea. Lee and Lawver (1995) used the pole of rotation to reconstruct the position of tectonic block of the Southeast China. Xia (2005) suggested that this tectonic pull apart by the rotation of Indochina should be supported by the mantle upwelling mentioned previously.

The tectonic evolution of the Southeast China and South China Sea began at the early Cenozoic (~60 Ma), at when Indochina was about 500 km NW of its present position. During the early Eocene (about 55-50 Ma), the northeastward moving Greater India began to collide with the southern margin of Eurasia and closed the Tethys Ocean. The collision of India and Eurasia forced the Indochina peninsula southeastwards along the left-lateral Ailao Shan-Red River fault and induced clockwise rotation of ~25° between 40 Ma and 22 Ma (Tapponnier et al., 1982, 1990, Lee and Lawver, 1995). The rotation of the Indochina peninsula led to the stretching and thinning of the continental shelf of South China block. The opening of the South China Sea started at the northwest part of the basin at ~32 Ma. The rheology of the South China block, Indochina and Sundaland blocks permitted the initial N-S directed extension of the South China Sea. The spreading direction changed during

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the early Miocene between magnetic anomaly 7 and 6b from predominately N-S to NW-SE direction. The directional change is thought to be related to ridge jump, which Barckhausen and Roeser (2004) suggests was due to the onset of a second spreading ridge in the southwestern part of the South China Sea basin however it remains uncertain (Hayes, 1988; Barckhausen and Roeser, 2004). The South China Sea expansion ceased at ~17 Ma due to the collision of the Kalimantan terrain and the North Palawan block (Holloway, 1982; Taylor and Hayes, 1983).

22 Figure 4. Map of Southeast Asia.

23 1.3.2 General Geology of Taiwan

The island of Taiwan is situated along the boundary of the Eurasian plate to the west and Philippine Sea plate to the east and is important location for understanding the geological and tectonic evolution of East and Southeast Asia.

Taiwan is ~385 km in length and ~143 km in width and has a total area of 35,960 km² and consists of six geological domains: Central Mountain Range, Tananao metamorphic complexes, Hsuehshan Range, Western foothills, Coastal Range and the Coastal Plain.

The collision of the Eurasia plate and Philippine Sea plate occurred during the Late Jurassic and created the central Mountain Range that reaches a maximum elevation of ~4000 meters. According to Seno (1977) and Suppe (1981), the collision rate of Philippine Sea plate with the Eurasian plate is about 7 cm/yr in a northwest-southeast direction (ω = 1.2°/ m.y., 45.5°N, 150.2°E) and causes the island to move

~3-4 cm/yr to the northwest and 1-2 cm/yr in vertical growth. Taiwan is located at the junction between the Ryukyu island arc and Luzon island arc thus there are characteristics of an island arc, accretionary wedge, continental plate and oceanic lithosphere. (Bowin et al. 1978; Chai, 1972; Suppe, 1984; Tsai, 1986; Kao et al., 1998)

Proto-Taiwan was formed during late Jurassic to early Cretaceous, about 150 million years ago, when the Eurasia plate and Pacific plate collided. The Pacific plate subducted westward beneath the Eurasia plate and created the Zhejiang-Fujian magmatic arc and proto-Taiwan. During the early Cenozoic, subduction slowed and eventually stopped allowing sediments to accumulate in the forearc basin. At the

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same time, the Philippine Sea Plate started to migrate northwestward and rotate clockwise from an E-W direct to its current N-S orientation (Lee and Lawver, 1995).

During the mid to late Miocene (~ 15-12Ma), the Luzon arc along a right-lateral transform fault pushes the continental shelf of the Eurasia plate whereas the other side of Philippine Sea Plate subducted under the extended Ryukyu arc (Teng, 1990; Huang et al., 2006). During the mid-Pliocene (~3-5 Ma), the continental shelf of Eurasia continually subducted under Luzon arc and consumed the accretionary wedge, which uplifted as a small island and rapidly produced a high mountain range (the proto-Central Mountain Range) during the late Pliocene (Teng, 1990; Huang et al., 2006). From the late Pliocene, the Luzon arc reached its current position off the coast of Hualien and the margin of the arc obliquely collided with Eurasia.

Concurrent with the Luzon-Eurasia collision, the tectonic setting of northeast Taiwan transformed from compression and collision to extension which caused the collapse of the Ilan terrain and the formation of the Ilan plain after a long period of orogenesis (Suppe, 1984; Lee and Wang, 1987; Teng, 1990, 1996, 2007). The volcanic rocks of the Luzon arc accreted to the uplifted Pliocene-Pleistocene passive margin sediments of Euraisa and became the Coastal Range whereas the uplifted sedimentary rocks developed into the Central Range. The northeast side of the Philippine Sea plate subducted beneath the Ryukyu Arc and induced backarc extension and the opening of the Okinawa Trough (Teng, 1990, 2007; Kao et al., 1998; Gao et al., 2008).

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Figure 5. Geologic Framework of Taiwan (revised from Teng, 1990).

26 1.3.3 Geological setting of Costal Range

The Coastal Range is located in eastern Taiwan and is one of the five main mountain ranges of the island. It stretches ~140 km north starting from the estuary of the Hualien River and ends at Beinan Mountain, Taitung. The Coastal Range is separated into a north part and a south part, bounded by Siouguluan River. The north part is slightly higher than the south part but is no longer being uplifted (Chen et al., 1991). River terraces are more abundant south of the Siouguluan River suggesting the area is still being uplifted. The Longitudinal Valley, between the Central Range and Coastal Range, is the boundary between Eurasia and the Philippine Sea plate.

The Central Range is underlain by deformed rocks of the Eurasia continental margin while the Coastal Range is underlain by island arc volcanic rocks from the Luzon arc (Page and Suppe, 1981).

Hsu (1956) divided the Costal Range into five formations by rock type, which from oldest to youngest are: Tuluanshan formation, Takangkou formation, Chimei volcanic complex formation, Lichi formation and Beinanshan conglomerate formation (Hsu, 1956, 1976). Biq (1969) suggested that the Takangkou formation and Chimei formation are a continuous sequence and named it after Takangkou formation. According to the petrographic differences, Teng (1979, 1980) added two new formations, the Baliwan formation and the Fanshuliao formation, into the original sequence described by Biq (1969). Teng and Wang (1981) proposed that Tuluanshan, Fanshuliao and Lichi formation belong to the pre-collision arc facies whereas the post-collision Baliwan and Beinanshan formations are continental facies.

In the paper by Teng and Lo (1985), they further classified the Tuluanshan formation

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as arc facies, the Takangkou formation as a fringing reef carbonate facies, the Fanshuliao formation as forearc basin deposition, and the Lichi formation as trench mixing facies. The Luzon arc was actively moving along a right-lateral transform fault during mid and late Miocene (~ 15-12 Ma) and collided with the continental margin during the mid-Pliocene (~ 5-3 Ma) followed by the formation of the Baliwan formation and Beinanshan conglomerate formation (Teng and Lo, 1985; Teng, 1990;

Huang et al., 2006).

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Figure 6. Geological setting of Coastal Range region (revised from Liou, 1979).

Figure 7. Different geological stratigraphy of the Coastal Range.

(Stratigraphy definition from Hsu, 1976; Chang, 1975; Chi et al., 1981; Teng and Wang, 1981; and Teng and Lo, 1985)

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1.4 East Taiwan Ophiolite and sample collection 1.4.1 Previous researches of the East Taiwan Ophiolite

The East Taiwan Ophiolite sporadically exposed along the southeastern part of the Coastal Range, from Taitung City to Hualien County, and considered to be a remnant of the South China Sea (Liou et al. 1977; Suppe and Liou, 1979; Liou, 1979;

Suppe et al., 1981; Jahn, 1986; Chung and Sun, 1992). The East Taiwan Ophiolite was first geologically mapped by Juan et al. (1953), and, at that time, was referred to as “Taiwanite” due to the large irregular shaped intrusion of glassy basalt (up to 95%

glassy texture). The geological mapping of the Coastal Range by Hsu (1956) outlined the full distribution of “Taiwanite”. Chen et al. (1976), Juan et al. (1976) and Chou et al. (1978) studied the mafic and ultramafic rocks and suggested that the “Taiwanite”

represented differentiated of a primitive melt. Chen et al. (1976) provided some mineral chemistry data and compared the crystallizing temperature of the olivine in

“Taiwanite”. Juan et al. (1976) performed the quenching experiment and EPMA, and discussed the characteristics of the alkali olivine basalt magma and olivine tholeiite magma from the “Taiwanite”. Chou et al. (1978) discussed the rare earth element pattern and ended up with the conclusion that the serpentinized peridotites may probably be the residua of mantle.

The book “The East Taiwan Ophiolite” written by Liou et al. in 1977 is an important progress of research of the ETO, in which they provided not only detailed reconnaissance record of the distribution and the rock sequence of the ETO but also new geochemical data and interpreting the ocean ridge origin of the ETO. Liou (1979) and Liou and Ernst (1979) were both discussing the metamorphism of the ETO and

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suggesting that the ETO experienced two stage of metamorphism, 1) pre-brecciation ridge-type metamorphism and 2) off-axis metamorphism. Suppe and Liou, (1979), Suppe et al. (1981), Jahn, (1986) and Chung and Sun, (1992) constrained the framework model of a mid-ocean ridge setting of the East Taiwan Ophiolite. Suppe and Liou (1979) summarized some previous studies (e.g. Page and Suppe, 1980;

Suppe et al. 1981) and suggested the concept of the tectonic model of the ETO. Suppe et al. (1981) reexamined the ophiolitic stratigraphy of the ETO and analyzed the serpentinized peridotite, gabbro, plagiogranite, diabase and basaltic rocks with XRF analysis. They also suggest the plausible mid-ocean ridge origination of the ETO and a displacement of 700-1000 km distance.

Jahn (1986) provided some new major and trace elements data, REE data and isotopic data and agreed with the ETO ocean-ridge interpretation. Radioisotopic age dating of the ETO and nanofossil studies (i.e. NN5) indicates that the ETO is ~14.6

± 0.4 Ma. Chung and Sun (1992) presented new major and trace elemental data, REE data and isotopic data and proposed that the ETO was formed at a normal slow-spreading axis environment. Shao et al. (2014) published new U-Pb age dates that shows the gabbroic rock of the ETO formed at ~17.5 Ma whereas the diorite and plagiogranite rocks were formed at ~14 Ma.

32 1.4.2 Sample collection

For this study samples were collected at Dianguang, Guanshan Township, along the Chia Wu Creek which is the largest exposed outcrop of the ETO. Rocks were collected along a 1.6 km traverse from midstream to upstream and ranged in size from pebble to boulder. The rock size generally increased from the midstream to upstream with some boulders as large as an automobile. The rock type also changed from midstream to upstream. The midstream rocks consisted of serpentinite, glassy basalt and pillow basalt outcroppings. Larger boulders of gabbro were found within the stream. Large boulders of gabbro and outcrops of peridotite were found near the headwater of the stream. Peridotite samples used in this research were collected around this spot. Subsequent to this fieldwork the outcrop was buried (summer, 2013) and cannot easily be reached.

Figure 8. Map of Guanshan area and Chiawu creek and the sample collecting spots.

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Picture 1. Serpentinite blocks.

Picture 2. Glassy basalt.

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Picture 3. Pillow basalt.

Picture 4. Gabbro.

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Picture 5. Fallen block of peridotite and landslide (behind).

Picture 6. Peridotite.

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Chapter 2

Research Methods

2.1 Petrography Description

Rock chips for thin sectioning were cut using a masonry saw to a size of ~3 x 2 cm. The pieces were pasted onto a glass slide (5 x 2.5 cm) with epoxy and placed on a hot plate for at least 24 hours until the epoxy solidified. The affixed sample was further cut to reduce thickness before polishing. The rock was polished using 200, 600 and 1000 mesh silicon carbide emery to grind the sample until proper thickness under microscope (~30 μm). Aluminum oxide powder (3 μm) was used for final polishing. An example of a finished thin section is found in picture 7.

A Carl Zeiss Axioplan 7082 polarizing optical microscope (Pic. 8) was used for the observation of 8 polished sections at the Department of Earth Science, National Taiwan Normal University. Individual minerals were identified using standard petrographic techniques (i.e. shape, color, cleavage, interference color, relief, extinction angle) and mineral modes were determined using a point-counter and visual estimation.

Picture 7. A finished thin section.

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Picture 8. Carl Zeiss Axioplan 7082 polarizing optical microscope.

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2.2 Scanning Electron Microscope (SEM)/ Energy Dispersive Spectrometry (EDS)

Small rock pieces (< 25 mm in diameter) were mounted in epoxy and subsequently ground and polish. The sample mounts were loaded into a JEOL USA JSM-7100F Analytical Field Emission SEM (Pic. 9) at the Laboratory of Electron Probe Micro-Analyses, Institute of Earth Science, Academia Sinica. The standard settings of this machine used for this study are: vacuum setting of the machine was set at 25 Pa, acceleration voltage at 15 kV, probe current at 0.18 nA, and working distance 10 mm. The focus, brightness and contrast of the field were adjusted for maximum clarity so that back scattered electron image pictures can be taken. The Oxford Instrument INCA-300 energy dispersive spectrometer (EDS) is attached on the SEM and was used to analyze the energy peak of targeted minerals. The duration of a single EDS analysis was 60 seconds, and the predicted mineral was judged depending on the energy magnitude of different elements.

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Figure 9. SEM structure.

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Picture 9. JEOL USA JSM-7100F Analytical Field Emission SEM.

43 2.3 Electron Probe Micro Analyzer (EPMA)

The electron probe micro analyzer (EPMA) was used to determine the major element composition of the targeted mineral. The EPMA detects the characteristic wavelength of different elements and quantifies the mineral composition. For the analyses, the sample mounts were coated with carbon using a Quorum Technology Q150TE high vacuum carbon coater and loaded into the EPMA. For this study, a JEOL W-EPMA model JXA-8900-R with 4 wavelength dispersive spectrometers (WDS) (Pic. 10) at the Laboratory of Electron Probe Micro-Analyses, Institute of Earth Science, Academia Sinica was used. The electron beam was defocused at an interval of approximately 10 mm on an area of about 5 µm diameter by beam conditions of 15 kV and 12 nA. Back-scattered electron images were used to guide the analysis on target positions of minerals. The measurements were corrected by using chemical-known standard minerals as listed: wollastonite for Si with TAP crystal and Ca with PET crystal, rutile for Ti (PET), corundum for Al (TAP), chrome oxide for Cr (PET), hematite for Fe with LiF crystal, tephroite for Mn (PET), pyrope for Mg (TAP), nickel olivine for Ni (LiF), albite for Na (TAP), and adularia for K (PET). Peak counting for each element and both upper and lower baselines are counted for 10 s and 5 s, respectively. Relative standard deviations (RSD) for all elements were less than 1%.

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Picture 10. JEOL W-EPMA JXA8900-R electron probe microanalyzer.

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2.4 In situ zircon Laser Ablation Inductively Coupled Plasma Mass Spectrometry (LA-ICP-MS)

Zircons were mechanically separated at the Yu-Neng Rock and Mineral Separation Co., Lanfang, Hubei province, China using a steel jaw crusher, magnetic separation and heavy liquids and purified by hand-picking under binocular microscope. Zircons were linearly mounted on a glass slide over a circle diameter of 1.5 cm. A mold was placed over the zircons and epoxy was poured over the zircons to ensure transfer of the minerals to the epoxy. The mount was polished to approximately half the mean grain thickness. A panchromatic CL imaging system (Gatan Mini-CL) attached to a scanning electron microscope (JOEL JSM-6360LV) was used to capture cathodoluminescence images (CL image) of the individual zircons to examine their internal structure. The epoxy zircon grain mount was loaded into a New Wave UP-213 laser attached to an Agilent 7500cx ICP-MS (Pic. 11 and 12) in the Tectonics and Low-temperature Heat Dating Laboratory, Department of Earth and Environmental Sciences, National Chung-Cheng University. Individual zircons were ablated using a beam diameter of 40 μm and frequency of 10 Hz. Helium gas was used as the carrier gas and supported by Argon gas. A single analysis takes 130 seconds, including 60 seconds preheating the tube and 70 seconds acquiring data.

The measurements were corrected by using age-known standards: GJ (600 ± 5 Ma) and Plešovice (337 ± 10 Ma).

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Picture 11. UP-213 laser ablation system.

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Picture 12. Agilent 7500cx ICP-MS.

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2.5 Thermal Ionization Mass Spectrometer (TIMS)

2.5 Thermal Ionization Mass Spectrometer (TIMS)