Petrological, geochemical, and hydrothermal characteristics of Ordovician cherts in the southeastern Tarim Basin, NW China, and constraints on the origin of cherts and Permian tectonic evolution

Accepted Manuscript Petrological, geochemical, and hydrothermal characteristics of Ordovician cherts in the southeastern Tarim Basin, NW China, and constraints on the origin of cherts and Permian tectonic evolution Jianhua He, Wenlong Ding, Wenhui Huang, Zicheng Cao, En Chen, Peng Dai, Yeqian Zhang PII: DOI: Reference:

S1367-9120(18)30452-8 https://doi.org/10.1016/j.jseaes.2018.10.030 JAES 3691

To appear in:

Journal of Asian Earth Sciences

Received Date: Revised Date: Accepted Date:

5 August 2017 6 October 2018 31 October 2018

Please cite this article as: He, J., Ding, W., Huang, W., Cao, Z., Chen, E., Dai, P., Zhang, Y., Petrological, geochemical, and hydrothermal characteristics of Ordovician cherts in the southeastern Tarim Basin, NW China, and constraints on the origin of cherts and Permian tectonic evolution, Journal of Asian Earth Sciences (2018), doi: https://doi.org/10.1016/j.jseaes.2018.10.030

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Petrological, geochemical, and hydrothermal characteristics of Ordovician cherts in the southeastern Tarim Basin, NW China, and constraints on the origin of cherts and Permian tectonic evolution ＊

Jianhua He a,b, Wenlong Ding a,b , Wenhui Huang a,b, Zicheng Caoc, En Chen a,b, Peng Daid, Yeqian Zhang e

a. School of Energy Resources，China University of Geosciences，Beijing 100083，China b. Key Laboratory for Marine Reservoir Evolution and Hydrocarbon Abundance Mechanism， Ministry of Education，China University of Geosciences，Beijing 100083，China c. Research Institute of Petroleum Exploration & Development, Northwest Oilfield Company, SINOPEC，Urumqi 830011，China d. Institute of Geophysical and Geochemical Exploration, Chinese Academy of Geological Sciences, Langfang 065000，China e. Guangdong Science & Technology Cooperation Center, Guangzhou 510030，China

* Corresponding author. E-mail address: [email protected] Postal address: School of Energy Resources, China University of Geosciences, Beijing 100083, China Tel.: +86 10 82320629 Fax: +86 10 82326850

ABSTRACT Cherts occur extensively as stratal wedges embedded in the Middle-Lower Ordovician carbonate successions in the circum-Tangguzibasi areas, south-east Tarim Basin, where thick (up to 2 km), dark grey well-stratified limestones deposited in shelf and platform environments are prominent. Based on detailed core and thin section observations, with petrographic, microthermometric, and geochemical data for the Middle-Lower Ordovician cherts in some key wells, two lithotypes of chert are identified across the marginal zone in the western Tangguzibasi areas: replacement and filled cherts. Replacement chert is characterized by granular microcrystalline quartz or radially fibrous chalcedony and commonly retains the pre-existing dolomite rhombs and elliptical ooids floating in these cherts. Filled chert, the most common type, is generally rimmed by fine-gained euhedral quartz that grades abruptly into vein-cavity-fill mega-quartz. Compositionally, these chert deposits are strongly enriched in Cr, Zn, Ba and U and generally show apparent right-convex REE patterns with low ∑REE values (avg. 4.24 ppm), weak negative Ce anomalies (Ce/Ce*: avg. 0.93) and low to intermediate positive Eu anomalies (Eu/Eu*: avg. 1.19), particularly the replacement cherts. Microthermometry of fluid inclusions from coarse quartz reveals high homogenization temperatures from 135 to 195 ℃ for trapped primary fluids with high salinities (11.0-18.9 wt%). These siliceous deposits also have low La/Ho and Th/U, relatively high Y/Ho and Ba/Sr, high Ce/Ce* (~1.06), intermediate to high (Lu/La)N (0.5-1.0) and low (La/Ce)N (mostly 1.31-1.45). All these data suggest that the cherts formed under the strong influence of hydrothermal fluids originating

from detritus-poor hydrogenetic Fe-Mn crusts at deep continental margins. Extensive silicification occurred after intense dolomitization and calcitization, while fracturing/faulting were greatly intensified in the host carbonates, induced by intense extension because of the uplift force of the magma cushion under oblique collision during the Early Permian. Most silica could have directly precipitated from hydrothermal fluids to chalcedony fibres around voids and vein cherts along fracture conduits. These petrological and geochemical data also provide a useful clue to better understand the extensive silica precipitation, intense tectono-hydrothermal activity and significant lithospheric fracturing events as a result of the closure of the South Tienshan and Palaeo-Tethys Oceans during the Early -Middle Permian. Keywords: Chert; Ordovician; hydrothermal activity; silicification; continental margin; Tarim Basin

1. Introduction Cherts containing up to 95% or more quartz have been widely studied in order to characterize palaeo-environmental changes (Eker et al., 2012; Ledevin et al., 2014), reconstruct biogeochemical/ocean conditions (Zhou and Xiao, 2007; McFadden et al., 2009; Xiao et al., 2010), and indicate tectonic setting and evolution (Chen et al., 2006; Wen et al., 2016). Cherts with various origins present different quartz cement features, sedimentary structures, and occurrence distributions, including mounded chert, vein chert, brecciated chert, bedded chert and nodular chert (Packard et al., 2001; Wang et al., 2012). Different types of chert have a wide range of origin models.

Recrystallization or redox diagenetic fronts are the general model for bedded or nodular chert formation as primary, early diagenetic cement in the upper-most sedimentary layers (Van den Boorn et al., 2007; Clayton et al., 1986), while vein or brecciated chert deposits result from Si- precipitation from hydrothermal and diagenetic fluids escaping into veins/fractures or replacement by Si- metasomatism of a sedimentary or volcanic protolith, respectively (Hofmannand Harris, 2008). The source of silica for chert remains problematic and complex. The silica for chert formation can derive from biogenic siliceous remains (e.g., radiolarians and spicules) (Boudreau, 1990), silica sediments inorganically precipitated from solution (e.g., siliceous sinter from hot spring waters and magadiite from saline waters) (Eker et al., 2012), volcanogenic silica related to submarine volcanisms (Adachi et al., 1986; Sugitani,1992; De Vries, 2004; Hofmann and Bolhar, 2007), directly Si-oversaturated seawater (Hesse, 1989; Sugitani et al.,1998; Perry and Lefticariu, 2003; Maliva et al.,2005; Tice and Lowe, 2006), or a mixture of both biogenic and inorganic sources (Frei and Polat, 2007; Van den Boorn et al., 2007, 2010; Marin-Carbonne et al., 2012). Over long geological times, chert layers may experience multiple -periods of silicification with various origins and natures of silica, so there is no unified model for interpretation of a certain chert. However, the criteria for the recognition of the various chert types provide a solution, including sedimentary structures (e.g., Maliva et al., 2005), petrological characteristics (e.g., Knauth and Lowe, 2003; Lu, 2012), and trace element or Rare Earth Elements (REEs) and/or isotope compositions (e.g., Kato and Nakamura, 2003; Bolhar et al., 2005; Van den Boorn et al., 2007, 2010;

Marin et al., 2010); the latter approach has become the most popular. In the Tarim Basin, the lower Palaeozoic contains multiple well-preserved chert beds, which are of great significance to reconstruct the palaeoenvironment and palaeoclimate conditions and to understand tectonic evolution. Moreover, these cherts are also primary sources of hydrocarbons (Feng and Liu, 2001; Yang et al., 2016) or directly become hydrocarbon reservoirs (Packard et al., 2001) and contain other economic mineral deposits (e.g., Mn, V and P) (Yu et al., 2004). Recently, various researchers have studied the depositional environments, petrology, sedimentary structures, geochemical characteristics and origin models of Cambrian cherts (Yu et al., 2004; Sun et al., 2004; Li et al., 2010; Chen et al., 2010; Yang et al., 2016). However, the geochemical and petrological characteristics and occurrence distribution of Ordovician cherts greatly differ from those of the Cambrian cherts (Jin et al., 2006), and little research has been conducted on the types, origin models and silica sources of the Ordovician cherts in the central-north part of the Tarim Basin (Li et al., 2015; Chen et al., 2016). Much less attention has been paid to the Ordovician cherts in the southern Tarim Basin, although different types of chert have been found in wells (Zhu et al., 2016). In this study, we focus on the Middle-Lower Ordovician cherts from some key wells in the southeastern Tarim Basin. Systematic petrographic evidence for these specific chert deposits, combined with microthermometric and geochemical data, is provided to (1) determine the petrology, trace element and REE characteristics of these cherts and the influence of hydrothermal alteration on pre-existing rocks; (2)

refine the origin of the local chert deposits and track the silica source; (3) reveal the diagenetic evolution of the Ordovician carbonate rocks and constrain the time of silicification; and (4) establish the trigger mechanism of silicification and form a better understanding of the tectonic evolution. 2. Geological setting 2.1. Sedimentary setting and stratigraphy The Tarim Basin, covering an area of approximately 5.6×105 km2, is the largest inland petroliferous basin in western China (Fig.1A). The area reaches the Tianshan tectonic belt in the north, the Altun tectonic belt in the southeast and the west Kunlun tectonic belt in the southwest. Based on the top surface of the basement (composed of crystalline Archean and Lower Proterozoic), the Tarim basin can be divided into three uplifts and four depressions, the Kuqa depression, the Tabei uplift belt, the North depression belt, the Central uplift belt (the Bachu Uplift, the Tazhong Uplift and the Gucheng nosed uplift), the Southwestern depression belt, the Tadong uplift belt and the Southeast depression from North to South (Jia, 1997) (Fig.1B). The Tangguzibasi Depression, extending in a nearly NE direction, lies in the eastern part of the Southwest Depression and covers about 3.65×104 km2 with an east-west length of 330-380 km and a south-north length of about 50-90 km. It is surrounded by the Bachu-Tazhong Uplift to the north, the South Uplift to the south, the Gucheng nosed uplift to the east, and the Maigaiti Slope to the west. The depression can also be divided into four tectonic sub-units, namely Yudong Fault belt, Madong Fault belt, Tanggu Sag and Tangnan Salient (Fig.1C). This study focuses on the Ordovician chert

successions across circum-Tangguzibasi areas, including the Yudong Fault belt, Madong Fault belt and Tangbei Fault belt. During the Cambrian and Ordovician, this area was dominated by carbonate and siliciclastic deposits from carbonate platforms to deepwater clastic basins (Fan et al., 2007). The Early Ordovician basin succeeded its depositional framework of the Late Cambrian. The carbonate platforms developed in the whole Tangguzibasi depression and expanded initially from southwestern part of the basin and the Tazhong Uplift. During this period, this area formed a set of restricted platform facies associations 300-500 m thick, consisting mainly of algal dolomite, dolostone and thinly bedded limestone (the Penglaiba Formation) and unconformably overlying the Upper Cambrian. Then, the Yingshan Formation is composed mainly of open platform deposites, with light gray or gray packstone, wackestone, and dolomitic limestone (0-223m) (Fig. 2A). However, the depositional framework of the Middle and Late Ordovician basin experienced significant changes and featured an open platform environment and widely developed reef/shoal-facies limestone along the platform margin and uplift (Gao et al., 2012), consisting mainly of thick-bedded packstone, grainstone, Sponge bafflestone and muddy limestone with a thickness of 150-700 m (The Yijianfang Formation, Qiaerbake and Lianglitage Formation). Then, Large-scale transgression occurred in the basin in the latter stage of the Late Ordovician (Lin et al., 2012). The carbonate platform was drowned and formed an extremely thick section of deepwater continental shelf- and slope-facies of dark mudstone and siltstone intercalated with calcarenaceous conglomerate and conglomeratic sandstone of

shelf-slope gravity and turbidite deposits overlying the carbonate platform(The Sangtamu Formation) (Fig. 2A). 2.2. Tectonic evolution and burial history During the Paleozoic, the Cambrian-Ordovician tectono-stratigraphic unit experienced different degrees of deformation, uplift and denudation in the various parts of the Tarim basin (He et al., 2016). In terms of the Ordovician successions in the circum-Tangguzibasi areas, there are three significant episodes of relatively active tectonic activity, including the late Early Paleozoic, the Middle Paleozoic and the late Permian, which have been reflected in the burial history (Fig. 3). During the late Early Paleozoic, a series of EN-trending thrust-faulted-and fault-bend folds (Fig. 1C and D) formed in the west-central part of the Tangguzibasi Depression similar to the South Altyn Fault in terms of their properties, phases and intensity (Tang et al., 2015). Simultaneously, the Penglaiba Formation and Yingshan Formation have been truncated by erosion because of the large-scale uplift movement in this area (Fig. 1D). During the Middle Paleozoic but before the Late Devonian, most regions of this area were uplifted so that the Ordovician strata from the margin to the more central basin were eroded step by step. These EN-trending faults formed in the Tangguzibasi Depression during the late Early Paleozoic were reactivated during the Middle Paleozoic, but the intensity and degree of faulting activity were significantly less than those in the latest Ordovician. During the late Permian, the tectonic activity in the Tangguzibasi Depression became weaker and most of faults stopped acting (Fig. 1C). However, before the late Permian, abnormal thermal event

also exerts a great influence on the Ordovician strata. During the Early-Middle Permian, the lithosphere of basin was broken up and large scale magmatism occurred while the uplift force exceeded the tortuosity of lithosphere.

The basic basalt lavas

and/or intrusive diabases, and intermediate acidic magmas distributed extensively in the central-northern part of basin, related to the evolution of Paleo-Tethys Ocean (Chen et al., 1997, 2006; Yang et al., 2005). In the circum-Tangguzibasi area, the Permian igneous rocks widely intruded into the Ordovician and its overlying strata in the western and northern. 3. Data set and methods 3.1. Data set In this study, core descriptions, cathodoluminescence imaging (CL), scanning electron microscopy (SEM) and thin section photographs, and samples have been obtained from the Ordovician carbonate from 8 cored wells in the southern slope of Tazhong Uplift and Yudong areas, with a total length of 153.83m and a current burial depth in the range of 3609.91 m to 6883.2 m (Figs. 1 and 2). In these samples, 48 samples were drilled from Yingshan Formation and 35 others were collected from Penglaiba Formation. Vertical plug samples are chosen from contrasting depositional and diagenetic environment to capture a wide range of silicification degree and origin in different types of carbonate rock and cherts. 3.2. Methods 3.2.1 Descriptive thin-section analysis, CL images and SEM observations The lithofacies and quartz texture were examined using an OLYMPUS

Polarizing Microscope at the China University of Geosciences (Beijing) from 10 polished thin sections of carbonates that were cut from core plugs. All of these sections were cut into 2 cm × 2 cm size. CL images were obtained with a CL8200-MK5 coldcathode instrument mounted on a petrographic microscope. A plane electron beam with a width of 3 mm is directed towards a polished thin section in a vacuum chamber attached to the microscope stage. The system was operated at 12 kV accelerating voltage and a current density of about 10 μA/mm2.This system is also fast, as no scanning is required, and can observe both colour and intensity of the CL. CL spectra were measured under standardized conditions (wavelength controlled by a Hg-halogen lamp, spot width 30 μm, measuring time 5 s). Multiple images were mosaicked to give a composite image that generally covered half a regular thin section. The

microporosity

characteristics

of

5

chert

samples

(cut

into

0.5cm×1cm×0.2mm chips) were observed using an “FEI Quanta model 200F” field-emission scanning electron microscopy (FE-SEM) with a working current set at 20kv and distance of 8-9 mm at the Microstructure Laboratory for Energy Materials of China University of Petroleum in Beijing. The composition of grains and cements was obtained by energy dispersive spectrometer (EDS) analysis to deduce the mineralogy. 3.2.2 Trace elements determination The 13 samples powders (40mg) for trace elements analysis were dried in an oven at 130 ℃ for 1-2 h and cooled to room temperature (20℃) at the CNNC Beijing

Research Institute of Uranium Geology, then dissolved in a tightly sealed Teflon screw-cap beaker with ultrapure 0.5 ml HNO 3+1ml HF+0.5ml HClO4 with a solubility of 8 mol/L and then dried. The dried samples were again digested with 1 ml 1% HNO3 +3ml H2O until a clear solution was obtained. The solution was diluted to 1:1000 by mass and analyzed on a Finnigan MAT inductively coupled plasma source mass spectrometer (ICP-MS) for measuring the trace and rare-earth element. Analytical precision for elemental tests was generally better than 4%. 3.2.4 Carbon and oxygen isotope analysis Carbon and oxygen isotope compositions of 36 samples were determined using a German Finnigan MAT 253 Isotope Ratio Mass Spectrometry (IRSM) at the CNNC Beijing Research Institute of Uranium Geology. The samples for the oxygen isotope measurement was first micro-drilled in the auto-carbonate mineral layers (deposited in the primary sedimentary environment), and the carbonate minerals associated with a hydrothermal chert front in the filled fractures or the vugs to know more information about the burial diagenetic process. The detailed isotopic measurements processes are referred as the standards of NBS19 and GBW04405. 3.2.5 Fluid inclusion measurement Fluid inclusion microthermometry and salinity of 19 samples were conducted using a LINKAM THMS600, Inc.-adapted, U.K. AIDI company-type, gas-flow heating-freezing stage mounted on an Olympus BX-53microscope. The samples were firstly mounted on a 0.23mm cover slip on a highly polished silver heating element to ensure excellent heat transfer and extremely sensitive temperature measurement. Then,

the samples will be placed on the instrument for heating. A platinum resistor sensor, accurate to 0.01°C provides far more accurate and stable temperature signal that can be achieved with a thermocouple. So the temperature can be stably controlled with heating rates of 0.01°C to 130 °C/min. Liquid-vapor homogenization temperatures were determined to ±0.05 °C by thermal cycling using temperature steps of 0.1 °C. To reduce the effect of re-equilibration on the homogenization temperatures (T h), the fluid inclusion assemblage (FIA), defined as a group of inclusions along a single growth zone with consistent vapour/liquid ratios (Goldstein & Reynolds, 1994), was applied in microthermometric measurement. The analytical accuracy of T h values and final melting temperature of ice (Tm) is within 0.5 ℃and 3 ℃, respectively. Salinity is calculated by Tm values and is expressed as wt% NaCl equivalent (Bodnar, 1993): wt% NaCl = 1.78×Tm – 0.0442×T2m + 0.000557×T3m.The experiments were conducted at a temperature of 26 °C and a relative humidity (RH) of 40%. Temperature range and trend records were correlated with the burial history curves for each area to specify the multi-periods of hydrothermal activity (e.g. silicification and dolomitisation) and also fracture opening and cementation. 4. Results 4.1 Petrology of chert The chert deposits occur as stratal wedges embedded in the carbonate successions (Fig. 2A). These cherts are restricted to the top and bottom sections of the Penglaiba Formation and the top of the Yingshan Formation. Two types of cherts (replacement chert and filled chert) can be further identified by occurrence

distribution, colour, micro-structure, quartz cement textures and origin models. 4.1.1 Replacement chert Replacement chert in cores macroscopically occurs as a grey-black, microcrystalline groundmass, distinguishable by its colour from the grey limestone and dark-brown dolostone (Fig. 4D). In thin section, crystalline silica appears in two forms. The first form, consisting of granular microcrystalline quartz, is determined to consist of euhedral ferroan dolomite and elliptical ooids floating in an extremely finely crystalline matrix (Fig. 4A and G). Most euhedral dolomite and elliptical ooids retain their original form and identity. It is clear that microcrystalline silica directly replaced the former host dolomite and limestone. For the most part, in replacement, chert occurs as diffuse masses in the contact zones between fine to medium crystalline dolomite, grainstone and partial packstone. A few microcrystalline grains form ctenoid shapes around the elliptical ooids (Fig. 4 G). The other form consists of isolated or intergrown fans of radially fibrous chalcedony (Fig. 4B). The fibrous chalcedony forms radiating fans or botryoids that are deposited in the former spaces or replace former minerals (e.g., calcite and dolomite). In scanning electron microscopy (SEM), radial silica characteristically contains abundant randomly oriented needle-like pseudomorphs with a common centre, subsequently replaced by silica (Fig. 4E). The needles have widths of 2-10 μm and are up to 100 μm or more long. The radiating micro-quartz cementation is similar to the texture of pillow basalt (Chen et al., 2010), which indicates a product precipitated from a rapidly cooling silica-rich fluid as a result of interactions between hydrothermal fluids and brines.

4.1.2 Filled chert Filled chert is more abundant than replacement chert in the Ordovician successions. The fibrous quartz locally appears in cores as either light grey to white or light brown fillings in former dissolution voids or cavities (Fig. 4F and Fig. 5D, J). This filling is also commonly rimmed by fine-gained euhedral quartz that grades abruptly into cavity-fill mega-quartz. In thin- section, two types of filled chert can be identified by the quartz cement texture and occurrence. Void chert, composed of chalcedony fibres, occurs as intergrown humps that form layers lining the walls of former cavities (Fig. 4C and H). The chalcedony fibres are composed of thin laminations identified by alternating light layers of fibrous quartz and/or dark thin laminae of fine-grained haematite (Fig. 5A). Locally, these fibres grade abruptly into long, slender, thicker, syntaxial quartz (Fig. 4H). The boundaries of quartz layers are defined by obvious breaks between zones containing different amounts of fine-grained haematite or pyrite that can be seen in plane-polarized light. In some samples, these laminations form beautifully zoned circular shapes that exhibit pseudo-uniaxial extinction under cross-polarised light (Fig. 4H and Fig. 5A). Chalcedony fibres are also indicators of silica polymer chain linearity, a function of pH levels of fluids (Heaney, 1993), which is favoured in rapidly cooling acidic siliceous solutions due to widespread suppression of silica polymer cross-branching (Hopkinson et al., 1999). The vein form, occurring as fracture cement, consists of highly euhedral, coarse-grained quartz (mega-quartz) or macro-quartz (Fig. 4C and G). In thin section,

macro-quartz crystals bend in opposite directions towards the fracture wall, indicating shear during antitaxial growth (Fig. 4C). Much fine-grained haematite can be found along the centre of the fracture (median line or zone). In SEM, the true nature of this macro-quartz is revealed as a mass of closely cemented subhedral to doubly terminated euhedral quartz prisms with lengths of 5-18 μm in the c-axis direction (Fig. 4I). This mega-quartz druse commonly crystallizes only after the silica concentration decreases, allowing slower nucleation and growth rates (Oethler, 1976). Therefore, the vein chert was probably precipitated along the early fracture/vein systems as a result of slowly cooling silica-rich hydrothermal fluids when they flowed upward, particularly as they flowed into the sub-branches away from the main fractures (Wang et al., 2012). Many small quartz crystals show primary precipitation and/or recrystallization during higher temperature, liquid-deficient episodes. Additionally, some micro-quartz fills the former voids or cavities and postdates at least one saddle dolomite precipitation event (Fig. 5B and F), replacing the former minerals along the fracture system (Fig. 4G). Thus, this observation further supports that silica-rich hydrothermal fluids from deep-seated hydrothermal sources migrated through the fault and/or fracture channels. 4.1.3 Hydrothermal minerals associated with silicification The mineral assemblage in a pre-existing rock can be altered to a new set of minerals that are more stable under certain hydrothermal conditions of temperature, pressure, and fluid composition and saturation. The interaction of hydrothermal fluid from depth and rocks forms a spatially and temporally regular zonal pattern of new

minerals as the cooling fluid moves through the surrounding rock mass along the fracture and cavity system (Inoue, 1995). Therefore, a certain mineral assemblage precipitated in a spatial and temporal order can provide obvious evidence of hydrothermal

activities

and

reflect

the

hydrothermal

environment.

Many

hydrothermal mineral assemblages are associated with silicification in the study area, including saddle dolomite, quartz, fluorite, pyrite, calcite, chlorite, anhydrite, barite and haematite. Saddle dolomite is the most common hydrothermal mineral and occurs as milky-white or caesious fillers in fractures or cavities (Fig. 5J). Saddle dolomites are generally close to faults and have filled in fractures (Fig. 5C). This observation indicates that the formation of saddle dolomites is related to fault-controlled fluid flow. In thin- section, the saddle dolomites consist of clear non-planar euhedral dolomite crystals with an average size of > 1 mm and are characterized by curved crystal faces (Fig. 5B and F). In some samples, fibrous quartz or fine-grained quartz cements cross-cut saddle dolomite crystals and precipitate in the inter-crystal pores between saddle dolomites. This relation further suggests that the micro-quartz precipitation event postdates the formation of saddle dolomite. Therefore, minor saddle dolomite crystals along some void and/or fracture channels could have been precipitated from the early hydrothermal fluids, with Mg2+ probably sourced from the Middle-Lower Cambrian dolostones through interactions with upwelling hydrothermal fluids. Fluorite filling often occurs as magenta minerals associated with coarse-grained calcite porphyroblasts and euhedral drusy quartz in dissolution-enlarged tectonic

fractures or solution cavities (Fig. 5D and G). Pyrite and barite crystals commonly appear as opaque minerals with fibrous quartz, wall-lining granular micro-quartz and drusy macro-quartz as a result of secondary heating and recrystallization in a relatively silica-poor fluid at a higher temperature (Alt et al., 1987). Therefore, the assemblage of amorphous silica, chalcedony fibres, euhedral mega-quartz and pyrite, as well as barite in some cases, can be attributed to variable super-saturation and rates of nucleation and growth in response to episodic hydrothermal activity (Hopkinson et al., 1999). 4.2 Trace element characteristics and REE geochemistry The trace element analysis results for cherts and their host rocks are given in Table 1. Ranges for some indices are also listed in Table 1. The upper continental crust (UCC)-normalised trace element patterns are shown in Fig. 6. Most chert samples in this study area are strongly enriched in Cr, Zn, Ba and U, whereas Rb, Hf, and Th show relatively less depletion than in the UCC (Fig. 6). Replacement cherts are characterized by high Ba (range: 262-389 ppm, avg. 323.5 ppm), Zn (range: 174-1638 ppm, avg. 585.5 ppm), and U (1.78-4.14 ppm, avg. 3.07 ppm) and relatively low Sr (range: 102-173 ppm, avg. 134.5 ppm) and Th (range: 0.303-0.557 ppm, avg. 0.409 ppm). Fracture-filling cherts also show relatively high Ba (avg. 27.36 ppm), Cu (avg. 121.89 ppm), and U (avg. 1.81 ppm), and low Th (avg. 0.235 ppm), Nb (avg. 0.253 ppm), and Ta (avg. 0.074 ppm) contents (Table 1). The concentrations of Zr, Cu, and Ni are consistent with a continental margin depositional environment (Murray, 1994). Generally, high Ba contents in some chert samples

reflect the high abundance of barite, reconciling a hydrothermal origin in view of the high Ba associated with hydrothermal fluids (Bertine & Keene, 1975; Urabe & Kusakabe, 1990; Halbach et al., 2002). Therefore, the concentration of Ba in silica deposits is a common characteristic of hydrothermal products. The Ba/Sr ratio can effectively distinguish between normal marine deposits (Ba/Sr <1) and hydrothermal deposits under the seafloor (Ba/Sr >1-104) (Peter and Scott, 1988), and characterise the degree of hydrothermal activity. Silica deposits with high Ba/Sr ratios are greatly influenced by hydrothermal processes (Smith and Cronan, 1983). Most of the samples in this study, except for a few samples of filled cherts, show high Ba/Sr ratios (avg. 1.38), particularly the replacement cherts (range: 1.87-3.67, avg. 2.04). This result suggests a strong influence of hydrothermal fluids on their precipitation (Table 1). The Th/U ratio in silica deposits can also reflect the degree of influence of Fe-Mg-rich materials derived from the lower crust or upper mantle (McLennan and Taylor, 1980; McLennan et al., 1990). Chert samples in this study area have extremely low Th/U ratios with a range of 0.06-0.26 (avg. 0.17), especially replacement cherts (avg. 0.15). This observation indicates that silica precipitation was strongly influenced by the Fe-Mg-rich hydrothermal fluid. Additionally, the Th/U ratio is generally used to distinguish the redox conditions (Wignall and Twitchett, 1996). For instance, wackestone has a high U content (2.93 ppm) and a low Th/U ratio (0.14), which reflects a primary anoxic marine environment (Table 1). The enrichment in U (avg. 2.44) and the low Th/U ratios in the chert samples also indicate silica precipitation under strongly anoxic diagenetic processes.

The chondrite-normalized and PAAS-normalized REE patterns for the chert samples are given in Fig. 7A and B. Generally, most of the samples exhibit apparent right-convex REE patterns, with weak negative Ce anomalies and low to intermediate positive Eu anomalies, particularly the replacement cherts, whereas a few samples of the filled cherts have slightly flat REE patterns with no obvious Eu anomalies. The ∑LREE/∑HREE ratios vary from 7.30 to 10.78 (avg. 8.60), which indicates enrichment in LREEs compared to HREEs in all types of cherts (Table 2). In a typical sea-floor hydrothermal system, REE patterns of vent fluids generally show an apparent depletion in HREEs relative to LREEs with pronounced positive Eu anomalies and weak or no negative Ce anomalies (Michard et al., 1983; German et al., 1990, 1999) (Fig. 7C) as a result of crysto-chemical exchange with plagioclase phenocrysts formed in the vent and speciation of vent fluids (Douville et al., 1999). Therefore, silica deposits derived from hydrothermal fluids generally inherit the REE patterns of the parent vent fluids, i.e., positive Eu anomalies (German et al., 1990; Mills and Elderfield, 1995), although REE patterns vary as the mixture of sea water increases. The REE patterns of the chert samples in this study are not similar to the typical sea floor hydrothermal system (Fig. 7 C and D). The total REE contents (∑REE) are very low in all types of chert (< 10 ppm), particularly in the filled chert (avg. 4.24 ppm) (Table 2). Eu/Eu* ratios from more than half of the samples are larger than 1 (avg. 1.2), particularly from the replacement cherts. This result also suggests a strong influence of hydrothermal fluids on silica deposition (Douville et al., 1999) (Fig. 8 A, B and F). Enrichment of Eu in cherts can

result from adsorption processes of various Ba-compound minerals (Dulski, 1994). The good correlations between positive Eu anomalies and Ba contents (R2 > 0.90) imply that the positive Eu anomalies are positively associated with increased Ba flux, which, in turn, is enhanced by hydrothermal activity during deposition (Fig. 8E). The lack of obviously positive Eu anomalies in fracture-filling chert suggests that this type of chert might have precipitated from the waning hydrothermal fluids near the end of hydrothermal activity and/or in the distal parts of fracture systems away from the main conduit zone (Wang et al., 2012). Except for one filled chert sample (YB3-W5), the Ce/Ce* ratios are less than 1 in all types of chert (avg. 0.93), representing weak negative Ce anomalies (Fig. 8B). In surface seawater with plentiful oxygen, Ce3+ released from particles is oxidized and then removed relative to the 3+ REEs, whereas in deep-sea sediments, Ce released from decomposing particles is less mobile than the 3+ REEs (Elderfield and Greaves, 1982). Therefore, the degree of Ce anomalies (Ceanom = lg [3CeN / (2LaN + NdN)]) is an effective indicator of the redox conditions, with Ceanom > -0.1 representing an anoxic environment and Ceanom < -0.1 representing an oxic environment. The Ceanom values of all types of chert are larger than -0.1 (avg. -0.05) (Table 2), suggesting silica deposition under an anoxic environment. The (La/Yb)N ratios range from 10.19 to 12.98 (avg. 11.63) in the filled chert and from 9.15 to 12.09 (avg. 10.18) in the replacement chert, showing enrichment in LREEs, especially for the filled chert samples (Fig. 8C). The (La/Ce)N ratios of all chert samples are less than 1.5 (avg. 1.39), which is consistent with continental margin deposition (Fig. 8D) (Murray, 1994). The value range of this ratio is quite different

from that of the Cambrian bedded cherts, which have a range of 2-3 in the northern Tarim Basin and were deposited on pelagic basin floors. 4.3 Carbon and oxygen isotopes Variations in the stable carbon and oxygen isotopes of carbonate rocks have been widely used for palaeoenvironmental reconstructions (Weissert, 1989, 2000). Carbon and oxygen isotope compositions of carbonate mineral cement fills in vugs and veins are also very important to reconstructions of diagenetic pathways in the study area because variations in carbon and oxygen isotope compositions are easily influenced by the fractionation effect of temperature, the chemical phase conditions and the surrounding rock environments (Godet et al., 2016; He et al., 2017). During burial diagenesis, different generations of cement or vein/vug-filling calcites or dolomites have been identified based on variable crystal size, CL colours, cross-cutting relationship between minerals, and homogenization temperatures. The host wackestone has relatively high δ13C values with a small range from 0.8 to 1.9 ‰ V-PDB (avg. 1.4 ‰) (Table 3; Fig. 9A), which is comparable to Ordovician seawater values (Gat, 1996) and indicates the primary depositional characteristics. The δ13C values of early calcite cement fills in the veins or vugs range from -0.2 to -1.8 ‰ V-PDB. These values are close to those of matrix dolomite and vug dolomite (between -0.8 and -1.2 ‰ V-PDB; Table 3; Fig. 9A). This result suggests an early stage of diagenesis under near-surface to shallow burial conditions in this study, including micritization, calcite cementation, meteoric water dissolution and dolomitization. In contrast, the late vein or vug calcite associated with silica deposits

(Fig. 5F) and saddle dolomite has a wider spread, ranging from -0.1 to -3.04 ‰ V-PDB, which shows a variable diagenetic environment. Additionally, compared to the change in δ13C values, δ18O values show a distinct range in each carbonate generation. Early calcite cements and dolomite show δ18O values from -4.2 to -8.7 ‰ V-PDB, while late diagenetic calcite cements associated with silica deposits and saddle dolomite show relatively lower δ18O values ranging from -8.3 to -14.3 ‰ V-PDB (Table 3; Fig. 9A). This change occurs because temperature during diagenetic processes (especially hydrothermal inputs) plays a much greater role in influencing the oxygen isotope composition, with a temperature increase of 1 decreasing the δ18O values by 0.24 ‰ (Talbot and Kelts, 1990). Previous researchers have established the quantitative oxygen- isotope fractionation equations for dolomite and calcite at various temperatures (Northrop and Clayton, 1966; Friedman and O’ Neil, 1977). High temperatures from hydrothermal fluids with strong thermal fractionation result in lower oxygen isotope values of calcite and dolomite, down to -14.3 ‰ V-PDB. The cross-plot of δ18O values and homogenization temperature can determine the δ18O values of the fluid source (Fig. 9 B and C). The δ18O values in the late calcite cements associated with silica deposits (4-8 ‰ SMOW) and saddle dolomite (1-3 ‰ SMOW) are significantly higher than those of sea water in the Cambrian and Ordovician (-8 -4 ‰ SMOW) and formation water (-4 – 1‰ SMOW). This result indicates that the formation of calcite-dolomite cement fills in vugs and veins is influenced by the external hydrothermal fluid. 4.4 Fluid inclusion microthermometry

Two-phase (liquid-vapour) primary aqueous FIAs are observed along growth zones of saddle dolomite and coarsely -crystalline vug/vein-lining quartz, calcite and barite crystals and are chosen to measure the homogenization temperature (T h) and final melting temperature (Tm). Some grains contain sparse, tiny inclusions that are invisible except at high- power, while others contain very abundant inclusions that appear as a ‘gritty’ or ‘dirty’ texture even at low power. Fluid inclusions are generally small (3 to 10 μm) and display a range of shapes (spherical, elongate or irregular), with vapour/fluid ratios of 4-8%. The individual FIAs have comparatively consistent Th and Tm values, indicating a minimum effect of re-equilibration (Goldstein, 2001). In this study, microthermometry was conducted on samples derived mostly from the Yingshan and partly from the Penglaiba Formation (Fig. 2A). Quartz crystals yield two Th peak values (Fig. 10A), indicating that the temperature of venting fluids in the same vent fields could be variable at different vents or venting episodes. The first peak value has relatively high Th values, ranging from 165 to 192 ℃ (avg. 173.5 ℃). The salinity, estimated from Tm values with a range of -7.8 to -10.8 ℃, varies from 11.0 to 14.7 wt% NaCl equivalent (avg. 12.9%). The second peak value has low Th values of 135 to 155 ℃ (avg. 143.8 ℃), Tm values between -11.5 and -15.3 ℃, and salinity from 15.5 to 18.9 wt% NaCl equivalent (avg. 17.8%) (Fig. 10). Calcite crystals also show a wider range of T h values from 81.6 to 190 ℃. The earliest vein and vuggy calcites have Th values from 81.6 to 118.7 ℃ (avg. 105.6 ℃), the lowest values in all cement mineral phases. Their Tm values range from -0.4 to -4.7 ℃, corresponding to salinity levels from 0.7 to 7.5 wt% NaCl equivalent (avg. 4.6%).

This generation of calcite was probably formed during the late Caledonian Orogeny (period of the fist oil charge: Jia et al., 2016). However, some of the late coarsemegacryst calcite (Fig. 5F) yield high T h values of 155 to 195 ℃ (avg. 163.5 ℃) and Tm values of -15.4 to -21.2 ℃ (avg. -18.0 ℃). The salinity calculated from Tm values varies from 19.1 to 23.5 wt% NaCl equivalent (avg. 18.0%). These values can be comparable to the Th values of typical hydrothermal minerals, such as saddle dolomite (range: 140 to 160 ℃) and barite crystals (avg. 193.3 ℃) (Fig. 10 A). Based on the assumption of an annual average surface- temperature of 20 ℃ and a normal geothermal gradient of 17-32 ℃/km (avg. 25 ℃/km) for the Middle-Lower Ordovician successions (Liu et al., 2017), the late calcite cements, saddle dolomite and barite crystals in these successions were likely triggered as these carbonates were buried to depths between approximately 5600 and 7500 m. However, considering the evidence for rather shallow burial depths for these rocks of approximately 4.2 to 5.5 km (Fig. 2B) and a maximum burial temperature of approximately 150 ℃ (Fig. 3), these extremely high fluid-inclusion temperatures probably result from hydrothermal activity. 5. Discussion 5.1 Constrains on silica source and depositional setting The occurrence of voids and/or fractures lined by macro- or mega-quartz and chalcedony cements in the Ordovician packstone and dolomite (Fig. 11C, E, and H) suggests that these features were formed along the fluid pathways generated through hydrothermal venting of silica-rich fluid migrating upward from depth but do not

represent an early diagenetic cement in the uppermost bedded sedimentary layers by primary chemical precipitation or biogenic siliceous remains. Fibrous silica cements with radiating fans or botryoidal structures are likely to have been a product precipitated from rapidly cooling silica-rich fluids as a result of interactions between hydrothermal fluids and formation brines (Herzig et al., 1988; Hopkinson et al., 1999). Chalcedony fibres with layers lining walls or zoned circular structures along the cavities or fracture (Fig. 5A; Fig. 11H and I) are favoured by a rapidly cooling acidic siliceous solution due to widespread suppression of silica polymer cross-branching as the indicators of silica polymer chain linearity, which is a function of the pH levels of fluids (Heaney, 1993; Hopkinson et al., 1999). On the other hand, the fibrous quartz or amorphous silica with replacement texture (Fig. 4A and G) was formed through silica permineralization or replacement of former carbonates, as these structures apparently are associated with the former carbonates (i.e., elliptical ooids and dolomite crystals). The presence of minor, wall-lining granular micro-quartz and drusy macro-quartz associated with barite and pyrite crystals in voids and veins (Fig. 4C; Fig. 11C and H) is probably a reflection of secondary mineral formation through secondary heating and recrystallisation in a relatively silica-poor fluid at a high temperature (Alt at al., 1987). Additionally, rare minerals, such as barite and fluorite in chert, are unlikely to be products of replacement by Si-metasomatism of carbonates or to occur through direct precipitation from normal sea-water. A more likely mechanism for introducing these minerals would be through the addition of hydrothermal inputs. For example, minor vug-filling barites (Fig. 5H) were probably

precipitated from upward flowing Ba-rich vent fluids upon mixing of relatively sulphate-rich formation waters with hydrothermal fluids (Bertine and Keene, 1975; Halbach et al., 2002). Hydrothermal sediments are relatively enriched in Cu and Ni and depleted in Co (Crerar et al., 1982). The high contents of Ba in silica sediments are commonly associated with hydrothermal fluids (Bertine and Keene, 1975; Halbach et al., 2002). Previous results (Table 1; Fig. 6) show that the Ordovician cherts are distinctively enriched in Ba, U, Zn and Cu followed by Cr and Ni to various degrees. The Co/Ni ratios of the cherts are much less than 1 (avg. 0.25), while the U/Th ratios are greater than 1 (avg. 7.56). In the Th-U diagram of various sediments, the chert samples all plot in a hydrothermal field (Fig. 12A), which is consistent with Amphitrite and Langban hydrothermal sediments. The origin model for these silica deposits is similar to that for Upper Cambrian and Ordovician replacement or vein-filled cherts on the northeastern slope of the Tazhong Uplift (Yu et al., 2004; Chen et al., 2010; Li et al., 2015). In hydrothermal processes, Cr is partly mobile, and it is enriched in hydrothermal precipitates without being followed by other “terrigenous” elements (e.g., Zr in zircon; Taylor and McLennan, 1985). In the Cr-Zr cross-plots (Fig. 12B), all the chert samples fall in the hydrothermal metalliferous sediment fields. The Th/U ratio in silica deposits can also reflect the degree of influence of Fe-Mg-rich materials derived from the lower crust or upper mantle (McLennan and Taylor, 1980; McLennan et al., 1990). Chert samples with extremely low Th/U ratios (avg. 0.17) indicate that silica precipitation was influenced by an Fe-Mg-rich hydrothermal fluid.

In medium-temperature (100-400 ℃) aqueous systems, high Y-Ho fractionation is common, as evidenced by hydrothermal vein minerals (e.g., fluorite or calcite and silica) (Bau and Dulski, 1995) that show non-chondritic Y/Ho ratios of up to > 200 (Bau, 1996). High Y/Ho ratios in chert samples (avg. 28.49) suggest a silica source from depth. The cross-plots of Th/U and Y/Ho ratios indicate siliceous precipitation from Si-saturated hydrothermal fluid driven by igneous activity (Fig. 12C). Additionally, the La/Ho ratios can identify the origin and migration of hydrothermal fluid (Bau, 1996; Chen et al., 2016). Hydrogenetic quartz vents have lower La/Ho ratios (<20) (Fig. 12D). As the silicification degree increases, La/Ho ratios sharply decrease, while Y/Ho ratios show little change (Chen et al., 2016). The low La/Ho and relatively high Y/Ho ratios of some replacement cherts (e.g., YB3-W2) suggest replacement by Si-metasomatic fluids from detritus-poor hydrogenetic Fe-Mn crusts compared with the filled cherts in the veins. REE abundance patterns and REE relative fractionations have proven useful environmental indicators in cherts of all ages and tectonic histories and important indicators identifying silica sources (Murray et al., 1990; Girty et al., 1993). Hydrothermal cherts are commonly characterized by total REE content, relatively right-inclined chondrite-normalized REE patterns and depletion in LREEs relative to HREEs with apparent positive Eu anomalies and weak or no Ce anomalies (German et al., 1990; Chen et al., 2006; Wang et al., 2012; Eker, 2012). Most chert samples show an apparent enrichment in LREEs relative to HREEs with intermediate positive Eu anomalies and weak negative Ce anomalies. Additionally, most of the studied

samples yield Pr/Pr* ratios <1.05 (avg. 1.01) (Table 2). This result indicates that the negative Ce anomalies are a result of La enrichment, rather than being authentically inherited from the sea-water characterized by depletion in LREEs relative to the HREEs with apparent Ce anomalies (Elderfield and Greaves, 1982). Unlike the hydrothermal-origin chert directly influenced by volcanic movements at the oceanic spreading centres (Herzig et al., 1988; Halbach et al., 2002), the silica source originated from depth immediately underneath the continental margin, rather than the hydrothermal fluids mixing with ancient seawater in the ocean. The Ce/Ce*, (La/Ce) N and (La/Lu)N ratios can further prove this interpretation. The Ce/Ce* values are very valuable for identifying the depositional environment of cherts in marine basins because of their stable values during diagenesis (Murray et al., 1992; Ding and Zhong, 1995). Chert sediments deposited on the oceanic spreading ridges have the lowest Ce/Ce* values (0.30±0.13), while those deposited on the ocean basin floor have intermediate Ce/Ce* values (0.60±0.13), and those deposited on continental margins have the highest Ce/Ce* values (1.09±0.25). The Ce/Ce* values of chert samples (0.87-1.06, avg. 0.93) are generally consistent with those of the continental margin environment, as supported above (Table 2 and Fig. 12E). This scenario is further demonstrated by the low (La/Ce)N values that mainly cluster between 1.25 and 1.48 (Fig. 12E), which are also in good agreement with the values of continental margins (0.5-1.5) proposed by Murray (1994). The intermediate to high (La/Lu)N values (0.60 to 1.03) fully support the scenario shown by Ce/Ce* and (La/Ce)N ratios and are also consistent with continental margins (Fig. 12F) (Murray et al., 1991). The cross-plots

of (La/Ce)N vs. Ce/Ce* and (La/Lu)N vs. (La/Ce)N show that Ordovician cherts were deposited in the continental margins (Fig. 12 E and F). Microthermometric data from fluid inclusions (Fig. 10A) indicate that the vug/vein-lining quartz and barite crystals in the chert samples were precipitated from hydrothermal brine fluids at temperatures of at least 135 ℃ (up to 195 ℃) within emanating channels (Fig. 10A), although temperature may be variable from channel centres to remote silicified layers. The high salinity of primary fluids (ca. three to five times sea-water salinity) indicates that the phase separation of magmatic-derived fluids must have occurred at higher temperatures at depth (Bodnar et al., 1985; Tivey et al., 1998). Additionally, hydrothermal megacryst calcite fills in the vugs/veins yield values of 145 to 200 ℃. The δ18O values of these calcite cements associated with silica deposits show relatively lower δ18O values ranging from -8.3 to -12.3 ‰ V-PDB (Fig. 9A and Fig. 10A). Previous researchers established the quantitative oxygenisotope (δ18O-SMOW) value equation for calcite at various temperatures (Northrop and Clayton, 1966; Friedman and O’ Neil, 1977), based on which the δ18O values of diagenetic fluids during the formation of hydrothermal megacryst calcite can be determined from +4.2 to +8.4 ‰ SMOW (Fig. 9B). Magmatic waters from depth typically show very uniform δ18O values with a range of +5.5 to +9.0. This similarity indicates that the high- δ18O values were acquired during direct or indirect interaction with magmatic waters after the magmas penetrated upward to shallow levels in the Earth’s crust. 5.2 Timing of silicification

Most replacement cherts and filled cherts incorporated or infiltrated into the Ordovician carbonate rocks clearly show that silicification of the dissolved host carbonate rocks by freshwater karstification occurred in the late Palaeozoic after deposition in the Ordovician. Silicification followed sediment consolidation and compaction, bulk dissolution (freshwater karstification), and much of the matrix dolomitization (Fig. 13). A degree of overlap with saddle dolomite cementation is suggested by the following. A large amount of internal sediments in dissolution cavities or fractures were dolomitized (Fig. 5 F and J; Fig. 11 E and G), but some of this sediment evidently infiltrated into the Penglaiba and Yingshan Formations after local matrix dolomitization. These limy internal sediments were then replaced by the silica deposits as the silicifying fluids moved through the fractures in the already dolomitized beds (Fig. 11 G). If the infiltration of these sediments into dissolution vugs and neptunian dykes (Fig. 5J; Fig. 11 E and F) was restricted in time, then the silicification is bracketed by two phases of matrix-replacive dolomitization. These cement saddle dolomites commonly fill or line dissolution vugs, moulds, and fractures (Fig. 5J; Fig. 11 B and D), indicating that their formation was associated with fracturing/faulting or tectonic events. Under these conditions, active faults and fractures are more likely to act as important pathways along which hydrothermal fluids from depth could have migrated upwards (Fig. 14) (Davies and Smith, 2006). Previous studies indicate that most faulting/fracturing activity occurred in these areas during the middle Palaeozoic but before the Late Devonian (Ren, 2011; He, 2016; Jia, 2016); then, these faults and fractures provided effective pathways for saddle

dolomite formation. Additionally, most vertical to subvertical chert veins cross-cut the saddle dolomites (Fig. 5 C and F; Fig. 11 C and E), and these saddle dolomites are considered to represent a precipitate of Late Devonian hydrothermal activity linked to collision events in the early Hercynian. This saddle dolomite phase has no hydrocarbons in primary inclusions and is therefore likely to predate Permian oil migration (Fig. 13). This interpretation also suggests that silicification was a post-Devonian event. The abnormal thermal events are mostly linked to hydrothermal activity for silica sources under regional extension. After the Devonian Period, the only abnormal thermal activity, the most intense activity, took place extensively in the study area and even across the entire basin during the Early Permian (Chen et al., 1997; Yang et al., 2005; Jin et al., 2006) and was characterized by basic basalt lavas and intrusive diabases and intermediate-acidic magmas. Therefore, the hydrothermal activity during the Early Permian was more likely the main silica source. In summary, the existing evidence suggests a Late Devonian to Early Permian age for chert emplacement. 5.3 Genetic models of silicification and basin infrastructure As such, the primary hydrothermal fluids undersaturated with carbonate could have caused dissolution of carbonate rocks along the fracture conduits during the course of upward migration (Fig. 14). The sharp lateral contacts between carbonates and chert deposits across the marginal zone further reveal the deep-seated, right-lateral strike-slip basement faulting along the pre-existing fault systems (Fig. 15), which also created a strong extensional regime in the distal part of this zone. The

syndepositional extensional faulting provided effective conduits for silica-rich fluids to migrate upward from depth, forming a silica inversion body on the basement. A small part of this microcrystalline silica directly replaced the former host dolomite and limestone (Fig. 14C). Furthermore, the dolomite crystals are best explained as a pre-existing phase set in a matrix that was subsequently silicified. The silicification process has retained the pre-existing dolomite rhombs floating in the chert (Fig. 4A), which indicates that replacement of the precursor substrate was on a microvolume basis. Most euhedral dolomite and elliptical ooids retain their pristine crystal faces (Fig. 4G) or have experienced relatively little micro-dissolution. This suggests that the H4SiO40-charged fluids were saturated or only marginally undersaturated with respect to dolomite (Gregg and Sibley, 1987). Therefore, the (ascending) silica-charged fluids would pass through dolomitized intervals with little reaction or potential buffering. This situation is interpreted as the genetic model for the replacement cherts. Another behaviour is that silica fluid could have oversaturated and directly precipitated to form chalcedony fibres around silica chimneys (or voids) and vein cherts along the fractured conduits as a result of conductive cooling and mixing with the deep brines in the carbonate intervals (Herzig et al., 1988) (Fig. 14D). Along fault/fracture and microfracture systems, these higher-temperature silica fluids migrated upward and laterally into the host carbonates, leading to hydrothermal recrystallization (or alteration) and dissolution of the previously formed matrix dolomites and subsequent precipitation of saddle dolomite cements (Fig. 14D). Then, hydrothermal fluids reached the top of the Ordovician and stopped further upward migration because of

the cap beds of the thick Carboniferous mud rocks (Fig. 14B). Therefore, unlike the hydrothermal-origin chert directly triggered by volcanic activity on the oceanic spreading centres (Herzig et al., 1988; Halbach et al., 2002), the filled chert formation is more likely interpreted as precipitation in the voids or veins from hydrothermal fluids fed by the magma inversion body immediately underneath. The regional variation in hydrothermal activity could also have been considered as

a

reflection

of

the

circum-Tangguzibasi areas.

intensity of This

basement

tectonic

activity

in

the

hydrothermal activity reflects an intense

magmatic-induced geothermal anomaly in the Central-Northern Tarim Basin during the Early Permian (Tian et al., 2010; Yang et al., 2007; Yu et al., 2011). Furthermore, the Early Permian was a crucial period during which the Tarim Basin environment was transformed from an epicontinental sea to an intracontinental lake. During the Cambrian-Ordovician, the Tarim Basin was a stable cratonic depression with marine basin, shelf and platform environments (Jia, 1997; Feng et al., 2005; Fan et al., 2007) (Fig. 15A). The sediments were dominated by carbonate rocks interbedded with minor mudstones in the Upper Ordovician. Before the Late Devonian, most of the Tarim Basin developed deep-water turbidities and terrigenous clastic and carbonate rocks on a continental shelf. Additionally, during the Silurian- Early Devonian stage, the most faulting/fracturing activity occurred in the study area, and then, the Ordovician strata from the margin to the more central basin were eroded step by step and developed a large number of dissolution pores (Fig. 15B). These pores could have provided important conduits for silicification. The Carboniferous sequence succeeded

the early depositional framework and was characterized by marine clastic rocks interbedded with carbonate rocks. However, during the Late Carboniferous-Early Permian, the sea completely receded due to extensive uplift; then, continental clastic rocks began to be widely distributed in this area. The South Tianshan oceanic crust was subducted beneath the Central Tienshan arc (Zhou et al., 2001; Xiao et al., 2004). The final collision of the northern Tarim block with the southern Siberian active continental margin is considered to have occurred in the Late Permian as a result of multiple, complicated amalgamation processes (Xiao et al., 2008). Additionally, the collision of the northern Tarim Basin with Palaeo-Tethys areas also occurred in the Early-Middle Permian (Yang et al., 2005; Xu, 2011). At the same time, the central-north-western part of the Tarim Basin experienced extensive uplift and intense extensional deformation because of the uplift force of the magma cushion and the broken lithosphere under oblique collision (Fig. 15 C and D). Accordingly, subsequent extensional strike-slip fault movement may also have taken place along pre-existing faults (Lin et al., 2015; Cunningham et al., 2003). Under this condition, these faults provided

favourable

conduit

systems

for

the

subsequent

migration

of

higher-temperature fluids from greater depths. The increased expelled fluids could have readily been driven to migrate upward by fierce heating and pressure loading along the extensional faults. These hydrothermal fluids then could have further migrated into the fractured Ordovician carbonate rocks and interacted with the host carbonate rocks above the basement, leading to silicification (replacement and precipitation) of carbonate rocks mainly along the fault/fracture networks (Fig. 15D).

6. Conclusions Based on detailed core and thin section investigations and petrographic, microthermometric, and geochemical studies on the Middle-Lower Ordovician cherts of some key wells in the circum-Tangguzibasi areas in the southern Tarim Basin, vital conclusions are drawn as follows: 1) Massive cherts occur as strata wedges embedded in the carbonate successions and are mainly restricted to the top and bottom sections of the Penglaiba Formation and the top of the Yingshan Formation. Two lithotypes of chert, including replacement chert and filled chert, can be further identified by their occurrence distribution, colours, micro-structures, quartz cement textures and origin models. 2) Most chert samples in this study area are strongly enriched in Cr, Zn, Ba and U, whereas Rb, Hf, and Th show lesser depletion. The high Ba/Sr ratios and low Th/U ratios for chert samples (especially replacement cherts) suggest a strong influence of hydrothermal fluids originating from the lower crust or upper mantle. The extremely low La (thereby ∑REE) abundances and right-convex REE patterns with weak negative Ce anomalies and low to intermediate positive Eu anomalies also point to a hydrothermal origin for the cherts, with significant REEs contributed from hydrothermal fluids. 3) The late diagenetic calcite cement associated with silica deposits and saddle dolomite with relatively lower δ18O values (mostly < -10 ‰) are obviously influenced by external hydrothermal fluid input. Most of the chert deposits were precipitated in hydrothermal brine fluids at higher temperatures of at least 135 ℃ (up to 195 ℃)

within emanating channels. The high salinities of primary fluids (ca. three to five times sea water salinity) indicate that the phase separation of magmatic-derived fluids must have occurred at higher temperatures at depth. 4）The high Ce/Ce* values (~1.06), intermediate to high (Lu/La)N (0.5-1.0) and low (La/Ce)N (mostly 1.31-1.45) in the siliceous deposits suggest that they were formed in continental margins. A silica source originating from detritus-poor hydrogenetic Fe-Mn crusts at deep continental margins can be further proved by the low La/Ho and relatively high Y/Ho and the Th-U diagram and Cr-Zr cross-plots. 5) Extensive silicification occurred while fracturing/faulting were greatly intensified in the host carbonates, induced by intense extensional deformation because of the uplift force of the magma cushion and the broken lithosphere under oblique collision during the Early Permian. Although part of the ascending silica-charged fluids replaced the former host dolomite and limestone, most silica directly precipitated from the fluids to form chalcedony fibres around voids and vein cherts along the fractured conduits as a result of conductive cooling and mixing with the deep brines in the carbonate layers. Acknowledgements： This research was supported jointly by National Natural Science Foundation Projects (GrantNos 41372139 and 41072098), National Science and Technology Major Project of China (2016ZX05046-003-001 and 2016ZX05034-004-003), Fundamental Research Funds for the Central Universities (2652017309) and China Scholarship Council. We are grateful to Exploration and Production Research Institute

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flood basalts from the Tarim Basin, Northwest China: SHRIMP zircon U-Pb dating and geochemical characteristics. Gondwana Res., 20, 485-497. Zhou, C., Xiao, S., 2007. Ediacaran δ13C chemostratigraphy of South China. Chem. Geol. 237, 89-108. Zhou, D., Graham, S.A., Chang, E., Wang, B. and Bradley, H. 2001. Paleozoic tectonic amalgamation of the Chinese Tian Shan: evidence from a transect along the Dushanzi-Kuqa Highway. In: Paleozoic and Mesozoic Tectonic Evolution of Central and Eastern Asia: From Continental Assembly to Intracontinental Deformation (Eds M.S. Hendrix and G.A. Davies), Geol. Soc. Am. Mem., 194, 23-46. Zhu, S.B., Wang, H.F., Wang, S.J., Zhou, Y., Zhao, Z.C., Wu, J.Q., 2016. Hydrothermal activities of Ordovician and its significance for alteration to carbonate reservoirs in Yubei area. Lithologic Reservoir. 28(3), 42-47.

List of Figures and Tables Fig. 1 Distribution and tectonic unit divisions of the Tarim Basin with the study area in the southeastern of Tarim Basin (A) and (B). The tectonic framework of the Tangguzibasi Depression showing the distribution of Paleozoic faults of the study area (the faults are projected to the top surface of Ordovician) (C). Seismic structure and Ordovician stratigraphic interpretation across Yudong Fault belt area (D). See Figure 1.C: A-A’ for location. Dash lines indicate the critical stratigraphic boundary; Faults are in their color definition. T75: Top of the early

Ordovician Yinshan Formation (O1+2y); T70: Top of the late Ordovician Sangtamu Formation (O3s); T56: Top of the early Carboniferous Systerm; T54: Top of the late Carboniferous System. Fig. 2 (A )The stratigraphic column of the Ordovician in the circum-Tangguzibasi areas of the Tarim Basin and lithological profile of the main sample wells; (B) Depth structure map on the top surface of Yingshan Formation (T 74) in the Yudong Area showing locations of the sample wells and also showing the distribution of Paleozoic faults. Fig. 3 Burial history of the Ordovician succession in the circum-Tangguzibasi areas of Tarim Basin based on the wells YB1 and YB5. Fig. 4 The core and thin-section characteristics and SEM photographs of the main types of the Ordovician cherts in the circum-Tangguzibasi areas. A. Well YB1-2, 5134.60m, O1-2y, cherty dolomite, the replacement microquartz postdates the formation of the saddle dolomite as these latter structures have been incorporated locally within the massive microquartz, retaining the original form and identity of dolomite crystals; B. YB3, 5360, O1-2y, limy replacement chert, showing intergrown fans of radially fibrous chalcedony infected with iron; C. YB5, 6841.8m, O1p, chert-bearing dolomite, a micro-vein filled by two generations of silica cements along the fracture wall in order: (i) thin microcrystalline bladed quartz crusts; (ii) thick meso-crystalline quartz mosaics; D. TB2, 4896.42m, O1p, grey-black replacement chert intruded into the dark-brown dolomite host, leading to progressive corrosion and/or replacement of dolostones. E. YB3,

5444.9m, O1p, chert-bearing wackestone, showing the radial silica with the needle shape. F. YB3, 5359.83, O1-2y, cherty packstone, showing the void/ or cave filled with the grey-black chert; G. YB 1-2, 5134.83m, O1-2y, limy replacement chert, showing the fracture and void system filled with fine-grain quartz and the ooids replaced with fibrous quartz, retaining their original shape; H. YB5, 6739.1, O1p, filled chert, the former dissolution cave filled with two types of cherts: (i) chalcedony fibres with obvious laminations forming beautifully zoned circular shape; (ii) long, slender, thicker, syntaxial quartz crystals; I. YB3, 5258.3m, O1-2y, chert-bearing dolomitic limestone, showing a mass of closely cemented subhedral to euhedral quartz prisms with a length of 5-18μm in the C-axis direction; (SD: saddle dolomite; MD: matrix dolomite; Ca: Calcite). Fig. 5 Typical core and thin-section pictures of the hydrothermal minerals in the Ordovician chert. A. YB5, 6739.1m, O1p, cherty dolomite, showing the laminated chalcedony fibres with alternating light layers of fibrous quartz and dark thin laminae of fine-grained hematite and pyrite; B. YB5, 6742.75m, O1p, cherty dolomite, showing the fibrous quartz cross-cut saddle dolomite cements; C. Z3, 4635.1m, O1-2y, wackestone, showing the saddle dolomite vein cross-cut the early calcite vein; D. YB1-2, 5444.05m, O1-2y, chert-bearing dolomite, showing the low-angle fracture filled with a mass of milky-white drusy macroquartz and the intergrown calcite and fluorite; E.YB3, 5263.5m, O1-2y, chert-bearing dolomitic packstone, showing the early dissolution cave filled with

masses of the golden pyrite crystals; F. YB5, 6743.23m, O1p, cherty dolomites, showing the calcite and coarse-grained quartz cements distribute between saddle dolomite crystals with wavy extinction, and cloudy centers surrounded by clear aggradational crystal growth under crosspolarized light; G. YB1, 5613.0m, O1-2y, chert-bearing packstone, showing the micrite replaced by fluorite and its intergrown calcite porphyroblast; H. YB1, 5598.0m, O1-2y, chert-bearing packstone, showing the coexistence of amorphous silica, barite and pirite in some cavities; I. YB1, 5564.47m, O1-2y, chert-bearing dolomite, showing the dolomite crustals replaced by anhydrite, associated with the amorphous silica and pyrite in some vugs; J. YB5, 6807.77-6808.37m, O1p, cherty dolomite, exhibiting extensive development of dissolution cavities filled with fibrous quartz and saddle dolomite; (MD: matrix dolomite; SD: saddle dolomite; Ba: barite; Ca: calcite; SI: silica; FI: fluorite; An: anhydrite). Fig. 6 Comparison of the enrichment factors of the trace elements (relative to upper continental crust, Wedepohl, 1995) in the cherts and host rocks. The dashed line represents an enrichment factor with a value of 1. Fig. 7 Rare earth element patterns (chondrite-normalized (A) and PAAS--normalized (B)) for cherts and host rocks from the Yudong region of Tarim Basin, which exhibits slightly negative Ce anomalies and HREE-depletion compared to LREEs. The chondrite-normalized average REE patterns of chert samples in the study area and Middle Cambrian chert from Li et al., 2010, compared with calculated patterns of mixed seawater and hydrothermal solution (C).

Chondrite-normalized average REE patterns of chert samples in the study area are plotted in D, compared with these of Lower Cambrian chert and Upper Permian igneous rocks. Chondrite REE data from Boynton, 1984; PAAS REE data from Taylor and McLennan (1985). Fig. 8 Cross-plot between Eu/Eu* and ∑LREE/∑HREE (A), Eu/Eu* and Ce/Ce* (B), (La/Ce)N and (La/Yb)N (C), Eu/Eu* and (La/Ce) (D), Eu/Eu* and Ba (E), U and Eu/Eu* (F). Fig. 9 Summary of isotopic and fluid inclusion data for principal nonchert mineral phase. (A) Relationship between δ18O (PDB) and δ13C (PDB) of principal nonchert mineral phase, including chert-bearing limestone, matrix dolomite, cherty dolomite filling in vugs, calcite cements filling in fractures and vugs and saddle dolomite; (B) Relationship between δ18O (PDB) of vein, vug calcite associated with silica deposits and homogenization temperature for various δ18O (SMOW) of fluid using fractionation equation by Friedman and O’ Neil, 1977, 103 lnαcalcite-fluid = 2.78×106T-2－2.89; (C) Relationship between δ18O (PDB) of vug dolomite and saddle dolomite and homogenization temperature for various δ18O (SMOW) of fluid using fractionation equation by Northrop and Clayton, 1966, 103lnαdolomite-fluid = 3.2×106T-2－3.3. Fig. 10 (A) Histogram of homogenization temperatures from primary fluid inclusions in key cement phases in the Ordovician chert and silicified host rocks. (B) Fluid inclusion fields for salinity and homogenization temperatures for primary inclusions.

Fig. 11 The core and thin-section characteristics and CL images of the Ordovician carbonate rocks and cherty carbonates, showing diagenetic sequences of different types of mineral phases. A. Well Z4, 3611.25m, O1-2y, wackestone, the late calcite with a dull luminescence under CL cross-cutting the early calcite with an orange color under CL; B. Well Z4, 5865.63m, O1p, dolomite, showing the wide fracture filled with subhedrel to anhedral saddle dolomite crystals of about 100 μm and up to 300 μm, displaying an orange-red color under CL, indicating that the pore fluid had a relatively high Mn/Fe ratios; C. Well YB5, 6742.75m, O1p, cherty dolomite, showing a micro-vein/channel filled by three generations of silica cements in order: (i) thin microcrystalline bladed quartz crusts; (ii) micro-crystalline to meso-crystalline quartz mosaics; (iii) cryptocrystalline fibrous chalcedony cements; D. Well Z4, 4285.73m, O1p, dolomite, showing the vug-fracture systems filled with three generations of cement calcite in order: (i) thick mesocrystalline-line calcite with a light orange color under CL; (ii) microcrystalline calcite cements with a dull luminescence; (iii) megacrysts calcite cements with an orange-red color under CL; enlarged dissolution pores also filled with saddle dolomite and ilmenite; E. Well YB3, 5371.52m, O1-2y, packstone, showing the fracture sealed with dolomite in the early stage, and then dissolute and filled with silica deposits in the center; F. Well YB3, 5359.83m, O1-2y, chert-bearing packstone, showing the enlarged dissolution pores along fracture system filled with the early megacrysts calcite, and then dissolute and filled with silica precipitation ; G. Well YB1-2X, 5445.1m, O1-2y, wackestone,

showing micro-crystalline calcite forms a lining along a fracture wall, surrounding saddle dolomite in the center; H. Well YB5, 6605.95m, O1p, cherty dolomite, showing a vein/channel in a microquartz host filled by two generation of silica cements: the early thin macro-megacrysts quartz forms a lining along a vein wall, inserted by the cryptocrystalline fibrous chalcedony cements in the center; I. Well YB5, 6739.1m, O1p, cherty dolomite, showing chalcedony fibres with gray-dark color under CL occurs as intergrown humps that form layers lining walls of former cavities. Fig. 12 Cross-plots of the trace elements and REEs for identification of various origins cherts. (A) Th vs. U from Rona, 1988;Ⅰ,TAG hydrothermal area;Ⅱ, East Pacific Rise crest deposits; Ⅲ, Red Sea hot brine deposits; Ⅳ, Galapagos spreading center deposits;Ⅴ, amphitrite hydrothermal sediments;Ⅵ, Langban hydrothermal sediments; Ⅶ, fossil hydrothermal deposits;Ⅷ, ordinary pelagic sediments;Ⅸ, ordinary manganese nodules; (B) Cr vs. Zr from Marchig, 1982;Ⅰ, hydrothermal metalliferous sediments;Ⅱ, diagenetic metalliferous sediments; Ⅲ, deep-sea sediments; (C) Y/Ho vs. Th/U; (D) La/Ho vs. Y/Ho; (E) (La/Ce)N vs. Ce/Ce*, the range of Ce/Ce* and (La/Ce)N from Murray, 1994; (F) (Lu/La)N vs. (La/Ce)N, the range of (Lu/La)N from Murray et al., 1991. *from Li et al., 2015; ** from Chen et al., 2010; ***from Yu et al., 2004. Fig.

13

Paragenetic

sequence

at

the

Middle-Lower

Ordovician

in

the

circum-Tangguzibasi areas. The temperature scale is based on depth and assumes a 20 ℃/km geothermal gradient. Solid time lines for diagenetic events represent

most likely timing; dashed lines reflect less probable or uncertain scenarios. The double-ended arrows under “magmatism” represent the ages (and error bars about those ages) for dated intrusive and extrusive materials that occur within the Tarim Basin along its central-western margin. Most dates are 40Ar/39Ar and K-Ar on basalt samples and are regarded as ages of emplacement. Magmatic bodies are mostly small-middle basic and alkaline intrusive. Data from Chen et al., 1997; Yang et al., 2005. A question mark indicates uncertain diagenetic events. Catagenesis data from Dong et al., 2013; Jia et al., 2016; Guo et al., 2016. Fig. 14 Conceptual trigger models showing the tectonic-stratigraphic texture (A and B) for the high temperature silicification (C), and also the development of hydrothermal venting fields and Ordovician chert deposition (D) on the circum-Tangguzibasi areas, Tarim Basin. See Figure 1.C: seismic profile for location. Fig. 15 Summary of block diagrams for the formation and evolution of the Ordovician chert in the Yudong areas, Tarim Basin, showing basin infrastructure and tectonic setting for chert formation in different key stages. (A) Early Ordovician period of regional uplift and meteoric water dissolution; (B) Early Devonian period of further regional uplift, thrust-fold faults and meteoric water dissolution, breccias and clastic deposits and minor hydrothermal dolomitization; (C) Early Permian period of extensional-wrench faulting, hydrothermal activity resulting in dissolution and dolomitization, and then silicification (emplacement of microporous replacement microquartz). (D) Distribution map of Permian igneous

rocks in Tarim Basin (modified by Yan et al., 2014) and also inferred regional and local stress fields, indicates most of Permian igneous rocks are distributed along the pre-existing basement faults, note the strain ellipse for the stress states and evolution. The central part of Tarim Basin (the blue rectangle) represents the most possible stress extensional field with an overall compressional tectonic regime. Table. 1 Concentrations of trace elements and Holmium of the Ordovician chert and its host rocks in the circum-Tangguzibasi areas Table. 2 Rare earth element results of the Ordovician chert and its host rocks in the circum-Tangguzibasi areas Table. 3 Stable carbon and oxygen isotope compositions results of wackestone and the host mineral filled in the vugs and veins of chert-bearing limestone and cherty dolomite

Figure1

Figure2

Figure3

Figure4

Figure5

Figure6

Figure7

Figure8

Figure9

Figure10

Figure11

Figure12

Figure13

Figure14

Figure15

Table. 1 Concentrations of trace elements and Holmium of the Ordovician chert and its host rocks in the circum-Tangguzibasi areas Elements

YB1-L1

YB1-L2

YB1-L3

YB1-L4

YB1-L5

YB3-W1

YB3-W2

YB3-W3

YB3-W4

YB3-W5

YB5-D1

YB5-D2

YB5-D3

UCC

Lithofacies

FC

FC

PC

FC

PC

RC

RC

RC

RC

FC

WK

LD

LD

/

Sc(ug/g)

0.947

0.846

1.07

1.06

1.17

0.747

1.14

1.51

1.58

1.26

0.915

0.593

0.673

16

V(ug/g)

7.04

6.26

5.87

6.15

7

12.1

6.74

10.9

10.2

18.3

19.3

17.6

19.5

98

Cr(ug/g)

7.41

9.68

5.98

6.37

6.61

466

407

33.2

24.9

40.2

11.3

16.3

25

126

Co(ug/g)

3.02

2.44

2.27

2.35

2.31

5.55

5.91

4.26

4.94

4.23

2.36

1.67

1.85

24

Ni(ug/g)

18.9

18.8

18.2

18.3

17.3

14.2

14.1

17

19.3

13.6

36

20.3

21.6

56

Cu(ug/g)

367

95.4

25.9

22.1

16.8

12

5.7

4.7

4.16

3.07

2.77

2.38

3.32

25

Zn(ug/g)

20.2

8.3

1.49

10.8

2.28

342

174

1638

188

304

6.08

6.62

3.75

65

Rb(ug/g)

1.98

2.35

1.69

1.76

2.09

2.28

1.42

5.31

2.23

1.49

2.8

1.34

1.78

78

Sr(ug/g)

114

119

120

117

124

106

102

157

173

168

223

142

96.9

333

Y(ug/g)

0.532

0.675

0.617

0.618

0.746

1.19

0.788

1.15

1.12

0.793

0.532

1.07

1.06

24

Ho(ug/g)

0.018

0.02

0.02

0.019

0.023

0.051

0.023

0.045

0.047

0.022

0.022

0.044

0.036

/

Zr(ug/g)

1.39

2.2

1.76

1.68

2.24

3.87

3.36

4.3

6.4

3.96

3.99

3.14

4.24

203

Nb(ug/g)

0.152

0.229

0.177

0.188

0.194

0.317

0.3

0.571

0.697

0.441

0.312

0.328

0.345

19

Cs(ug/g)

0.122

0.15

0.105

0.114

0.125

0.12

0.085

0.203

0.119

0.055

0.103

0.054

0.092

3.4

Ba(ug/g)

8.64

69.9

51.7

15.7

362

389

262

293

350

15.2

20.1

34.4

23.3

584

Hf(ug/g)

0.041

0.059

0.046

0.047

0.053

0.097

0.085

0.129

0.123

0.151

0.092

0.132

0.113

4.9

Ta(ug/g)

0.017

0.021

0.016

0.021

0.018

0.038

0.068

0.102

0.081

0.043

0.029

0.117

0.129

1.1

Th(ug/g)

0.17

0.267

0.22

0.205

0.257

0.334

0.303

0.557

0.443

0.297

0.419

0.372

0.453

8.5

U(ug/g)

0.785

1.05

0.798

0.768

0.756

4.14

3.15

3.19

1.78

4.65

2.93

0.629

0.572

1.7

Th/U

0.217

0.254

0.276

0.267

0.340

0.081

0.096

0.175

0.249

0.064

0.143

0.591

0.792

/

Ba/Sr

0.076

0.587

0.431

0.134

2.919

3.670

2.569

1.866

2.023

0.090

0.090

0.242

0.240

/

Y/Ho

29.556

33.750

30.850

32.526

32.435

23.333

34.261

25.556

23.830

36.045

24.182

24.318

29.444

/

Note: lithofacies, RC: Replacement chert; FC: Filled chert; PC: Packstone; WK: Wackestone; LD: limy dolomite; UCC data from Wedepohl, 1995

Table. 2 Rare earth element results of the Ordovician chert and its host rocks in the circum-Tangguzibasi areas Elements

YB1-L1

YB1-L2

YB1-L3

YB1-L4

YB1-L5

YB3-W1

YB3-W2

YB3-W3

YB3-W4

YB3-W5

YB5-D1

YB5-D2

YB5-D3

Chondrite

PAAS

NASC

Lithofacies

FC

FC

PC

FC

PC

RC

RC

RC

RC

FC

WK

LD

LD

/

/

/

La(ug/g)

0.77

0.887

0.82

0.801

0.9

1.66

1.12

2.26

1.94

1.22

0.781

1.33

1.35

0.31

38

32

Ce(ug/g)

1.48

1.73

1.64

1.53

1.79

2.91

2.03

3.99

3.52

2.53

1.59

2.65

2.94

0.808

79.6

73

Pr(ug/g)

0.169

0.185

0.182

0.176

0.204

0.369

0.261

0.454

0.421

0.237

0.152

0.254

0.317

0.122

8.83

7.9

Nd(ug/g)

0.595

0.714

0.683

0.701

0.807

1.63

0.975

1.74

1.58

0.863

0.606

1.07

1.06

0.5

34

33

Sm(ug/g)

0.095

0.14

0.106

0.117

0.137

0.362

0.2

0.318

0.268

0.144

0.15

0.214

0.258

0.195

5.55

5.7

Eu(ug/g)

0.025

0.048

0.036

0.032

0.029

0.158

0.097

0.14

0.137

0.035

0.047

0.046

0.049

0.0735

1.08

1.24

Gd(ug/g)

0.097

0.161

0.137

0.121

0.174

0.217

0.194

0.22

0.268

0.144

0.076

0.145

0.238

0.259

4.66

5.2

Tb(ug/g)

0.018

0.021

0.019

0.019

0.024

0.051

0.03

0.039

0.033

0.024

0.016

0.025

0.038

0.0474

0.77

0.85

Dy(ug/g)

0.096

0.115

0.114

0.106

0.131

0.282

0.15

0.215

0.285

0.189

0.098

0.178

0.189

0.322

4.68

5.8

Ho(ug/g)

0.018

0.02

0.02

0.019

0.023

0.051

0.023

0.045

0.047

0.022

0.022

0.044

0.036

0.0718

0.99

1.04

Er(ug/g)

0.053

0.062

0.063

0.059

0.078

0.119

0.097

0.132

0.137

0.09

0.048

0.089

0.185

0.21

2.85

3.4

Tm(ug/g)

0.008

0.009

0.008

0.009

0.008

0.019

0.015

0.019

0.011

0.01

0.007

0.018

0.021

0.0324

0.41

0.5

Yb(ug/g)

0.04

0.054

0.051

0.053

0.065

0.114

0.078

0.126

0.143

0.067

0.059

0.106

0.17

0.209

2.82

3.1

Lu(ug/g)

0.006

0.009

0.009

0.008

0.009

0.015

0.015

0.017

0.028

0.019

0.011

0.016

0.015

0.0322

0.43

0.48

∑REE

3.470

4.155

3.888

3.751

4.379

7.957

5.311

9.715

8.818

5.597

3.663

6.185

6.866

/

/

/

∑L/HREE

9.253

8.106

8.150

8.439

7.496

7.985

7.303

10.777

8.119

8.839

9.730

8.886

6.642

/

/

/

Eu/Eu*

0.789

0.975

0.913

0.816

0.574

1.596

1.487

1.536

1.547

0.799

1.204

0.755

0.595

/

/

/

Ce/Ce*

0.947

0.978

0.981

0.941

0.968

0.860

0.874

0.897

0.897

1.065

1.045

1.029

1.047

/

/

/

Pr/Pr*

1.03

0.98

1.01

1.00

0.99

0.98

1.06

1.02

1.05

0.93

0.91

0.88

1.05

Ceanom

-0.037

-0.035

-0.028

-0.055

-0.040

-0.107

-0.077

-0.073

-0.068

0.007

-0.012

-0.026

0.016

/

/

/

(La/Ce)N

1.356

1.336

1.303

1.365

1.311

1.487

1.438

1.476

1.437

1.257

1.280

1.308

1.197

/

/

/

(La/Yb)N

12.978

11.074

10.840

10.189

9.335

9.817

9.681

12.093

9.146

12.276

8.924

8.459

5.354

/

/

/

Note: lithofacies, RC: Replacement chert; FC: Filled chert; PC: Packstone; WK: Wackestone; LD: limy dolomite; Eu/Eu* = EuN / (0.5SmN + 0.5GdN); Ce/Ce* = CeN / (0.5LaN + 0.5PrN); Pr/Pr* = PrN / (0.5CeN + 0.5NdN); Ceanom = lg [3CeN / (2LaN + NdN)], CeN, LaN and NdN normalized by North America Shale Composite (NASC); Chondrite REE data from Boynton, 1984; PAAS REE data from Taylor and McLennan (1985); NASC data from Gromet et al., 1984

Table. 3 Stable carbon and oxygen isotope compositions results of wackestone and the host mineral filled in the vugs and veins of chert-bearing limestone and cherty dolomite Well

YB1-2

YB3

Host Mineral

Samples

δ13C V-PDB (‰)

δ18O V-PDB (‰)

δ18O V-SMOW (‰)

Host Mineral

Samples

δ13C V-PDB (‰)

δ18O V-PDB (‰)

δ18O V-SMOW (‰)

VEC

YB1-2-C1

-0.3

-4.3

26.5

VD

YB3-C17

-1.4

-8.7

22

MD

YB1-2-C2

-1.2

-6

24.8

VD

YB3-C18

-0.8

-8.4

22.2

VD

YB1-2-C3

-1

-7.5

23.2

VD

YB3-C19

-1.3

-7.9

22.7

VUCS

YB3-C20

-3

-10.2

20.4

well

YB3

VUCS

YB3-C1

-1.98

-10.83

19.7

VEC

YB3-C2

-1.2

-5

25.7

VUCS

YB3-C21

-1.8

-10.2

20.4

VUC

YB3-C3

-0.6

-5.2

25.6

VUCS

YB3-C22

-1.8

-10

20.5

VEC

YB3-C4

-0.2

-4.2

26.6

VUCS

YB3-C23

-0.4

-9.6

21

VUC

YB3-C5

-0.8

-4.9

25.8

MD

YB5-C1

-1

-6.3

24.4

VEC

YB3-C6

-0.4

-5

25.8

SD

YB5-C2

-2.2

-12.12

17.85

VUC

YB3-C7

-1.7

-5.4

25.4

SD

YB5-C3

-1.3

-11.3

19.2

VUC

YB3-C8

-1.8

-4.6

26.1

SD

YB5-C4

-3.04

-14.3

16.11

MD

YB3-C9

-2.5

-7

23.6

VECS

YB5-C5

-2.35

-11.86

18.64

MD

YB3-C10

-0.9

-5.4

25.3

VUCS

YB5-C6

-2.32

-11.62

18.88

MD

YB3-C11

-1.2

-6.1

24.6

VUCS

YB5-C7

-1.82

-11.09

19.43

SD

YB3-C12

-1.6

-11.3

19.3

WC

YB9-C1

0.8

-7.4

23.3

VECS

YB3-C13

-0.1

-8.3

22.3

WC

YB9-C2

1.8

-3.6

27.2

VECS

YB3-C14

-0.6

-8.8

21.8

WC

YB9-C3

1.9

-3.3

27.5

VECS

YB3-C15

-0.8

-8.4

22.3

WC

YB9-C4

1.2

-4.5

26.3

VECS

YB3-C16

-1.5

-9.5

21.1

YB5

YB9

Note: Host mineral: VEC: vein-calcite; VUC: vug-calcite; VECS: vein-calcite associated with silica deposites; VUCS: vug-calcite associated with silica deposites; WC: wackestone; MD: matrix dolomite; SD: saddle dolomite; VD: vug dolomite

Highlights    

Replacement and filled cherts of the Ordovician were identified by occurrence distribution, sedimentary structures and geochemistry characteristics. Silica deposits originated from detritus-poor hydrogenetic Fe-Mg crusts at deep continental margins after late Devonian Trigger mechanism of silification was controlled by tectono-hydrothermal activity. Silification represents significant lithospheric fracturing events during the Early -Middle Permian.

Graphical abstract