Application of Stable Strontium Isotope Geochemistry and Fluid Inclusion Microthermometry to Studies of Dolomitization of the Deeply Buried Cambrian Carbonate Successions in West-Central Tarim Basin, NW China

0. INTRODUCTION

The Tarim Basin in Northwest China is the largest oil-bearing basin in China (Fig. 1). The carbonate rocks in the Tarim Basin account for more than 50% of oil and gas production, with a greater percentage of hydrocarbon reserves in carbonate produced from dolomites due to their relatively high porosity and permeability. Dolomites occur pervasively in the West-central Tarim Basin and are up to 1 692 m thick (Hu et al., 2009; Wu et al., 2008; Shao et al., 2002; Hu and Jia, 1991). The carbonate rocks in the study area have undergone complex diagenetic history from early shallow marine to deep-burial environments. Carbonate reservoirs related to tectonics have been the topic of interest in the Tarim Basin in recent years (Han et al., 2017; Xu et al., 2015).

Figure  1.  A schematic tectonic map of the Tarim Basin, showing the distribution of structural units within the basin and the location of wells in the study area (after Lin et al., 2009).

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Numerous hydrological models have been proposed to explain the widespread occurrence of dolomitized rocks, particularly for deep-burial dolomites (Davies and Smith, 2006; Machel, 2004). The mechanism of dolomitization and the nature of dolomitizing fluids have been a mystery to geologists for centuries despite a series of models of dolomitization that has been put forward to interpret the ubiquitous occurrences of dolomite in stratigraphic records (e.g., Han et al., 2016; Guo et al., 2016; Jiang et al., 2016, 2015; Dong et al., 2013a; Tucker and Wright, 2009). Recently, hydrothermal dolomitization has drawn much attention from researchers than ever before and is becoming a new research focus in the Tarim Basin (Guo et al., 2016; Jiang et al., 2015; Dong et al., 2013a; Chen et al., 2004). Because dolomites from different origins usually have distinct petrographic and geochemical characteristics, geochemical studies have been used to characterize the composition of the dolomitization fluids to fully understand the dolomitization processes (Guo et al., 2016; Jiang et al., 2016; Zhu et al., 2015; Dong et al., 2013a, b; Veizer et al., 1999). Although the studies carried out in this Basin on the Lower Paleozoic carbonates, which has been interpreted as a product of multiple mechanisms of dolomitization (Guo et al., 2016; Jiang et al., 2016; Dong et al., 2013a), no consensus has been reached on the genesis and geochemical studies of the deeply buried Cambrian dolomites in the west-central parts of the basin.

This paper will delve into the dolomitization of the deeply buried Cambrian carbonate successions in the West-central Tarim Basin. A multidisciplinary approach is used which combines petrography and geochemistry (O-C-Sr-isotopes), and fluid inclusion microthermometry in order to (ⅰ) elucidate the dolomitization of deeply buried Cambrian carbonates and (ⅱ) formulate a conceptual model to elucidate the source of dolomitizing fluids and regimes of fluid flow in the study area.

1. GEOLOGICAL SETTING

This Tarim Basin has experienced multiple stages of tectonic evolutions, including the Caledonian, Hercynian, Indosinian and Himalayan cycles (He et al., 2016; Tang, 1997), which has affected both the early basin structure and later sedimentation. Two episodes of tectonic processes occurred in the early stage of the Caledonian tectonic movement, with the first episode being the Keping tectonics occurring at the end of the Sinian and resulted in the unconformity (T₉⁰) between Upper Sinian basement rocks and the Lower Cambrian strata (Fig. 2a), developing a series of extensional faults during the early stage of the Caledonian tectonic movement (He et al., 2016). The mid-Caledonian tectonic movement started at the end of Cambrian and continued to Early Ordovician, forming a regional unconformity (T₈⁰) between the Upper Cambrian and Lower Ordovician successions, and later transformed cratonic basin into a foreland basin (Gao and Fan, 2014). During the Lower and Middle Ordovician Period, the basin and its adjacent areas developed a stable cratonic depocenter, with carbonate platforms developing in its northern, central and southwestern parts, and sloped to deep sea environment in the eastern part (Gao and Fan, 2014). The carbonate platform gradually became restricted to the west and a central part of the basin (Fig. 2a) until it was completely flooded in response to a marine transgression (Gao and Fan, 2014).

Figure  2.  (a) A stratigraphy cross-section showing various sedimentary facies of the Cambrian strata in the West-central Tarim Basin (modified from Jiang et al., 2016); (b) burial history curves of the Cambrian strata, an integral modification of wells H4 and Z4 in the West-central Tarim Basin (modified from Qiu et al. 2012). Note the estimated major diagenetic event (dashed-green rectangle) and the dashed-red rectangle indicates the timing of cement phase precipitation during the Early–Late Permian Period.

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The study area is located in the central uplift belt which is oriented west-to-east of the Tarim Basin, and comprises the Bachu uplift in the west and Tazhong uplift in the center with high-dip thrust faults (Fig. 1) (He et al., 2016; Gao and Fan, 2014). The Sinian rested unconformably on the pre-Sinian crystallized basement and composed of siliciclastics and phosphorite deposits, etc. The Cambrian can be divided into three sub-sequences as follows: the Lower Cambrian, which includes the Yuertusi, Xiaoerbulake and Wusongger formations; the Middle Cambrian includes the Shayilike and Awatage formations; and the Upper Cambrian consists of the Xiaqiulitage Formation (Fig. 2a). the Lower Cambrian consists of siliciclastics rocks, thick dolomites, limestone and evaporites. The Middle Cambrian contains thick gypsum, anhydrite and salt layers, dolomites, limestones, and the Upper Cambrian consists of thick dolomites, limestones etc., characterized by restricted to evaporative-lagoonal carbonate platform and open platform environment (Fig. 2a) (Gao and Fan, 2014).

The burial and geothermal history for the Cambrian successions in the West-central Tarim Basin is illustrated in Fig. 2b. This history shows that rapid sedimentation took place at the phase of the passive continental margin and the Cambrian strata were quickly buried to a depth of about 5 000 to 8 000 m in the northwest of Tarim Basin before it continued subsiding to its current depth (Fig. 2b; Jiang et al., 2016; Qiu et al., 2012). The Tarim Basin has been subjected to four major episodes of abnormally thermal events since the Middle–Late Neoproterozoic, Cambrian–Early Ordovician, Permian and Cretaceous periods (Chen et al., 1997). The thermal event that took place during the Early Ordovician was localized in the central-western part of the basin (Chen et al., 1997). The most intense thermal event, which occurred at 290.5±2.9 Ma as constrained by U-Pb isotopic dating (Dong et al., 2013b), occurred extensively in the basin during the Early Permian and was characterized by magmatic activities mostly in the central and northern parts of the basin (Chen et al., 1997).

2. METHODS

About 180 core samples of Cambrian carbonate were collected from eight recently drilled exploration wells in the West-central (Bachu-Tazhong) Tarim Basin (Fig. 1). Representative samples were selected for detailed petrolgraphic description, ¹³C, ¹⁸O and ⁸⁷Sr/⁸⁶Sr isotopes analyses and fluid inclusion microthermometry.

Detailed petrographic study was conducted on 110 polished thin sections from dolomites, vug-fracture-filled dolomite and calcite cement from different specimens collected at different intervals from bottom to top in eight well-cores in the Cambrian strata. Some of the thin sections were prepared from dyed, resin-impregnated samples and some were half stained with Alizarin-Red-S (Dickson, 1966) to differentiate calcite from dolomite and ferroan carbonates from nonferroan carbonates, and they were examined by petrographic microscope in the Key Laboratory of Exploration Technologies for Oil and Gas Resources, Yangtze University. Cathodoluminescence (CL) microscopy was performed on a RELIOTRON Ⅲ stage from RELION Industries (Bedford, MA, USA) with a 5 to 8 kV beam and a current intensity of 300 to 500 μA, at the Institute of Geology and Geophysics, Chinese Academy of Sciences (IGGCAS). Fluorescence microscopy was employed using an Olympus BX51 microscope (Olympus Corporation, Tokyo, Japan) with a 365 nm wavelength excitation using a high-pressure mercury lamp at IGGCAS.

Microthermometry measurement of fluid inclusions was performed on selected doubly polished thin sections (60 mm thick) on a Linkam THM600 heating-cooling stage at IGGCAS. Calibration with precision of ±1 ℃ at 300 ℃ and ±0.1 ℃ at -56.6 ℃ was conducted using synthetic H₂O and CO₂ fluid inclusion standards. Heating experiments were conducted prior to conducting cooling experiments so as to reduce the risk of stretching or decrepitating the fluid inclusions. The accuracy of homogenization temperature (T_h) and final melting temperature of ice (T_m) values are within 2.5 and 0.5 ℃, respectively. Salinity levels in weight percent (wt.%) were calculated from final melting temperature (T_m) using the equation of Bodnar (1993) in terms of the H₂O-NaCl system: wt.%NaCl=1.78×T_m–0.044 2×T_m²+0.000 557×T_m³.

Representative samples from dolomites, limestones and vug-fracture-filled dolomite and calcite cement from different specimens were extracted mechanically from clean rock surfaces using a tungsten-tipped dental drill for ¹³C and ¹⁸O isotope analysis at the IGGCAS. Approximately 20 mg powder of the samples was reacted with anhydrous phosphoric acid (H₃PO₄) at 25 ℃ for 24 and 72 h, respectively. The analysis of the produced CO₂-gases released from the sample was conducted with a Finnigan MAT-252 mass spectrometer. The correction of oxygen isotope values of the samples was carried out using the fractionation factor of 1.011 78 (Rosenbaum and Sheppard, 1986). The carbon (δ¹³C) and oxygen (δ¹⁸O) values were expressed relative to the Vienna Peedee Belemnite (VPDB) standard. Precision of measurement was monitored through routine analysis of NBS-19 standard (δ¹³C=+1.95‰ VPDB; δ¹⁸O= -2.20‰ VPDB). The standard deviation for both δ¹³C and δ¹⁸O values was better than ±0.1‰ (1δ).

For strontium (⁸⁷Sr/⁸⁶Sr) isotope analysis, powdered samples of 60 mg from same selected representative samples for ¹³C and ¹⁸O isotope analysis were dissolved in 2.5N HCL on the hot plate at 90 ℃. The final sample solution was loaded into a chromatographic column with 2 mL of AG50Wx12 200-400 mesh cation exchange resin for separation of strontium from the sample matrix. Isotopic measurement of isolated Sr was performed on a Finnigan MAT-262 thermal ionization mass spectrometer (Thermo Fisher Scientific, Waltham, MA, USA) at IGGCAS, and corrected for instrumental mass fractionation according to exponential law with ⁸⁸Sr/⁸⁶Sr of 8.375 209. All ⁸⁷Sr/⁸⁶Sr values were normalized to a NBS-987 ratio of 0.710 253. Analytical precision (2σ) was monitored and controlled by repetitive analyses of international standard sample NBS-987. The mean standard error (2σ) is ±12×10^-6.

3. RESULTS

3.1 Petrography

Three main types of crystalline dolomite textures were recognized, based on crystal size and crystal boundary shape (planar and non-planar) (Sibley and Gregg, 1987), including very fine- to fine-crystalline dolomite (RD1), fine- to medium-crystalline dolomite (RD2) and medium- to coarse-crystalline dolomite (RD3). Only the dominant medium- to coarse-crystalline dolomite (saddle) cement (CD) is described.

3.1.1   Crystalline dolomites

3.1.1.1   Very fine- to fine-crystalline, nonplanar-a to planar-s, dolomite (RD1)

The RD1 is a nonplanar-a to planar-s dolomite varying in crystal size from 20–60 μm (Figs. 3a–3e), showing well-preserved precursor depositional texture and occurs in some instances as "clustered or floating" subhedral-anhedral crystals in matrix along low-amplitude stylolite in grainstone (Figs. 3a, 3b). It shows somewhat interfingering of rhomb-like dolomite crystals with cloudy cores and clear rims. Under CL, the RD1 crystals display a dull-red luminescence in the cores and bright-red thin rims (Figs. 3b and 3d). In volume, the RD1 accounts for less than 1% of the dolomite rocks in this study and are common in the Shayilike Formation and Xiaqiulitage Formation.

Figure  3.  (a) Photomicrograph of RD1 showing scattered subhedral "clustering or floating rhomb-like" crystals in matrix along a low-amplitude stylolite in grainstone, well ZS1S8, 5 820.4 m, under PPL; (b) CL photomicrograph of (a) showing dull red luminescence and CD dolomite partially filling a dissolution vug; (c) RD1 showing nonplanar-a to planar-s textures, with fracture-filled CD dolomite cement, well TZ1S16, 6 429.7 m, under PPL; (d) CL photomicrograph of (c) showing dull red luminescence, with fracture-filled CD dolomite cement displaying a dull red luminescence in the cores with light red rims; (e) RD1 showing fracture-filled with early-stage calcite cement (red arrows) and bitumen (Bt), well TZ3S10, 5 436.2 m, under PPL; (f) RD2 showing crystalline pyrite disseminated within intercrystalline pores (red circles), well TIS15, 4 479.8 m, under PPL; (g) RD2 showing planar-e(s) dolomite crystals displaying turbid cores with clear rims, well Z4S9, 6 721.2 m, under PPL; (h) CL photomicrograph of (g) showing dull red luminescence in the cores and bright red thin rims; (i) RD2 showing planar-s textures with fracture-filled later-stage calcite cement (LCAL), well ZS5S6, 6 124.6 m, under PPL; (j) CL photomicrograph of (i) showing dull red luminescence, with fracture-filled LCAL displaying a light orange luminescence; (k) RD3 showing planar-e dolomite crystals, with early-stage calcite cement (ECAL) filling dissolved vug and CD dolomite cement lining fractures, well H4S17, 5 727.4 m, under PPL; (l) CL photomicrograph of (k) showing dull red luminescence, with ECAL displaying dark orange reflection; (m) RD3 with dolomite cement (CD) growing into vugs filled with acicular-bladed gypsum crystals (GYS), well Z4S12, 7 236.5 m under XPL; (n) RD3 showing anhedral dolomite crystals, with fracture-filled ECAL (yellow arrows) cross-cut by LCAL (red arrows), well F1S18, 4 412.6 m, under PPL; (o) RD3 with CD dolomite cement crystals displaying turbid cores with clear rims growing into empty vugs, well T1S16, 4 544.7 m; (p) CL photomicrograph of (o) with CD dolomite cement showing dull red luminescence in the cores with light red rims.

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3.1.1.2   Fine- to medium-crystalline, planar-e(s), dolomite (RD2)

In hand specimen, RD2 is grey to dark grey in colour. It is characterized by planar to rectilinear crystal surfaces and well-defined crystal boundaries with turbid cores and clear rims (Figs. 3f–3j) and 60–250 μm in size. Some samples show crystalline pyrites disseminated within intercrystalline pores (Fig. 3f). Intercrystalline pores and vugs are presented and some remain open forming the major type of porosity (Figs. 3g–3j). Some samples have their dissolved vugs and fractures filled with early and later-stage calcite cement (Figs. 3i–3j). Their crystals show a sharp extinction under cross-polarized light, and display a dull-red luminescence in cores and bright-red thin rims under CL (Figs. 3h and 3j). The RD2 constitute about 8%–10% by volume of all dolomite rocks and are sparsely common throughout the Cambrian strata in the study area.

3.1.1.3   Medium- to coarse-crystalline, nonplanar-a to planar-s, dolomite (RD3)

In hand specimen, RD3 is grey to light grey in colour. It exhibits recrystallization related fabrics. They usually display rhombohedral to saddle-like shaped crystals with cloudy cores and clear rims (Figs. 3k–3m). Their crystal sizes vary from 50 to 500 µm, with somewhat curved and tightly packed crystal faces especially in those samples close to fault/fracture zones (Figs. 3m–3p). Intercrystalline pores, dissolution vugs and fractures are the main types of porosity and are open or partially or completely occluded by calcite cement, CD dolomite and acicular-bladed gypsum (Figs. 3k–3p). They display undulose extinction under cross-polarized light. Under CL, they show purple to red luminescence in cores and dull-to bright-red rims (Figs. 3i and 3p). They occur mainly in the Lower and Upper Cambrian strata. The RD3 is the most abundant dolomite and constitutes roughly 80% of the total dolomites in the study area by volume.

3.1.1.4   Medium- to coarse-crystalline, nonplanar-a, dolomite cement (or saddle dolomite) (CD)

The CD (saddle) dolomite cement is creamy white or pink in hand specimen and range in crystal size from 250 μm–3 mm (Figs. 3c, 3k, 3m and 3o). Their crystals are characterized by mildly curved bladed or lobate-like crystal faces with half-moon (or scimitar-like) crystal terminations (Figs. 3c, 3l and 3m). It occurs exclusively as cement infilling or lining of serrated edges along fractures and/or dissolved vugs in crystalline dolomites, and partially or completely occludes the vugs or fractures (Figs. 3b, 3c, 3k, 3m and 3p). Under CL, they usually exhibit dull-red luminescence in the core and bright-red or light-orange rims (Figs. 3b, 3d, 3i and 3p). Volumetrically, the CD dolomites constitute about 2%–5% of all dolomites.

3.1.2   Later stage diagenetic minerals

As the precipitation of dolomite cement ceases, later-stage calcite cements, gypsum cement and pyrite were precipitated as infills in residual dissolved vuggy-pores, intercrystalline pores and fracture (Figs. 3f, 3j, 3l, 3m and 3n). The calcite cement can be divided into two stages based on their cross-cutting relationships (Fig. 3n). The early-stage calcite cement (Figs. 3e, 3l and 3n) is commonest vug and fracture filling and the later-stage calcite cement (Figs. 3j and 3n) is mostly fracture filling. On the other hand, the sulphide mineralization consists of pyrite and gypsum. Pyrite occurs within intercrystalline pores (Fig. 3f). Acicular-bladed gypsum crystals occurred as pore-fillings (Fig. 3m). Volumetrically, the later stage diagenetic minerals constitute less than 1% of the mineral phases in the dolomite rocks.

3.2 Isotope Geochemistry

3.2.1   Oxygen and carbon isotope

The carbon and oxygen isotopic compositions of the analyzed samples are summarized in Table 1 and plotted in Fig. 4a. The limestone samples have δ¹⁸O and δ¹³C values ranging from -9.9‰ to -7.3‰ VPDB (average -8.7‰, n=10) and -1.3‰ to -0.3‰ VPDB (average -0.7‰), and are within the range reported for carbonates in equilibrium with Cambrian seawater (δ¹⁸O: -10‰ to -6.0‰, and δ¹³C: -2.5‰ to 1.5‰; e.g., Shao et al., 2002; Veizer et al., 1999).

Table  1.  Isotope data (δ¹³C, δ¹⁸O and⁸⁷Sr/⁸⁶Sr) of limestones (LST), RD1-RD3, CD dolomite cement, and early and later-stage calcite cement in the Cambrian carbonate successions, West-central Tarim Basin

Sample code	Lithology	Formation	Depth (m)	δ18O (‰, VPDB)	δ13C (‰, VPDB)	87Sr/86Sr	(±2σ)
H4S13	LST	Є2a	5 256.6	-9.4	-1.2	0.709 216	±10
F1S6	LST	Є3x	2 696.4	-8.6	-0.3
F1S9	LST	Є3x	3 728.6	-7.9	-0.6	0.709 025	±10
F1S16	LST	Є2s	4 125.3	-9.6	-0.4	0.708 971	±12
H4S15	LST	Є2s	5 362.5	-7.8	-0.7	0.708 879	±12
H4S18	LST	Є1x	5 733.4	-9.9	-1.3
H4S10	LST	Є2a	5 246.7	-7.6	-0.7	0.709 247	±11
T1S8	LST	Є1w	4 356.4	-7.3	-1.1
H4S14	LST	Є2s	5 346.3	-8.9	-0.3
F1S11	LST	Є3x	3 742.8	-7.5	-0.5
TZ3S12	RD1	Є3x	5 445.7	-9.4	-0.9	0.709 331	±10
ZS1S6	RD1	Є3x	5 711.6	-9.6	-0.7
Z4S5	RD1	Є3x	6 452.8	-8.6	-0.7	0.709 276	±11
TZ3S10	RD1	Є3x	5 436.2	-8.3	-0.6	0.709 261	±10
TZ3S13	RD1	Є2a	5 716.7	-8.4	-0.4
ZS1S8	RD1	Є3x	5 820.4	-9.8	-1.3	0.708 975	±12
ZS5S4	RD1	Є3x	5 844.3	-9.6	-1.2	0.709 221	±10
ZS5S5	RD1	Є3x	5 865.6	-8.2	-0.8	0.709 327	±11
TZ1S16	RD1	Є3x	6 419.4	-8.9	-1.4	0.709 312	±10
TZ3S14	RD1	Є2a	5 729.2	-7.1	-0.4
Z4S7	RD2	Є3x	6 453.5	-9.9	-3.6	0.709 279	±11
H4S8	RD2	Є3x	4 934.4	-10.9	-3.2	0.709 025	±10
T1S15	RD2	Є1w	4 479.8	-8.4	-2.4
Z4S9	RD2	Є2a	6 721.2	-7.9	-2.8	0.709 386	±10
H4S16	RD2	Є2s	5 448.6	-8.7	-3.6	0.709 089	±11
TZ1S20	RD2	Є3x	6 429.7	-8.9	-3.3	0.709 414	±12
H4S12	RD2	Є2a	5 252.3	-8.6	-3.1	0.709 047	±11
TZ3S15	RD2	Є2s	5 812.4	-7.1	-0.5	0.709 287	±10
ZS1S10	RD2	Є2a	6 223.3	-9.2	-3.4	0.708 876	±11
ZS1S11	RD2	Є2a	6 236.5	-9.7	-3.9
ZS5S6	RD2	Є3x	6 124.6	-10.3	-3.7	0.709 374	±12
ZS5S7	RD2	Є2a	6 235.4	-8.7	-2.5
T1S12	RD2	Є2s	4 212.7	-8.6	-2.7	0.709 256	±10
Z4S12	RD3	Є1w	7 236.5	-7.6	-3.6	0.709 271	±11
Z4S13	RD3	Є1x	7 733.8	-10.4	-3.1	0.709 341	±11
ZS1S14	RD3	Є2a	6 420.4	-8.4	-3.5	0.709 235	±10
TZ3S18	RD3	Є2s	5 853.6	-7.8	-2.9
T1S16	RD3	Є1x	4 544.7	-8.4	-0.4	0.709 329	±10
T1S17	RD3	Є1x	4 557.3	-9.6	-3.1	0.709 322	±09
ZS1S15	RD3	Є2s	6 521.6	-10.1	-3.4	0.708 992	±10
H4S17	RD3	Є1x	5 727.4	-10.2	-3.3
F1S18	RD3	Є1w	4 412.6	-9.5	-3.9	0.709 421	±12
F1S19	RD3	Є1x	4 480.3	-8.3	-0.5	0.709 402	±11
H4S19	RD3	Є1x	5 745.5	-9.8	-2.7	0.709 113	±10
ZS5S8	RD3	Є2a	6 334.7	-9.4	-2.5
ZS5S10	RD3	Є2s	6 419.4	-9.8	-3.2
ZS5S12	RD3	Є2s	6 515.2	-10.3	-3.4
ZS5S14	RD3	Є2s	6 530.6	-9.6	-3.7
ZS1S17	RD3	Є1w	6 724.3	-8.7	-2.8
T1S18	Vein(CD)	Є1x	4 562.5	-9.1	-3.7	0.709 234	±10
ZS1S10	Vug (CD)	Є2a	6 223.3	-9.6	-3.6	0.708 983	±12
ZS5S6	Vein (CD)	Є3x	6 124.6	-8.4	-3.2
TZ1S12	Vug (CD)	Є3x	5 819.5	-7.8	-3.4	0.709 265	±11
ZS1S14	Vein (CD)	Є2a	6 420.4	-8.8	-3.5
TZ3S12	Vug (CD)	Є3x	5 445.7	-7.9	-1.5
F1S19	Vein (CD)	Є1x	4 480.3	-8.3	-1.9	0.709 321	±10
T1S6	Vug (CD)	Є2s	4 231.7	-10.6	-2.7	0.709 271	±11
Z4S12	Vein (CD)	Є1w	7 236.5	-10.2	-2.9
TZ1S20	Vein (CD)	Є3x	6 429.7	-10.3	-3.4	0.709 351	±12
T1S16	Vein (CD)	Є1x	4 544.7	-10.1	-3.2
Z4S13	Vein (CD)	Є1x	7 733.8	-9.8	-3.5	0.709 412	±11
T1S17	Vug (CD)	Є1x	4 557.3	-10.4	-2.7	0.708 964	±10
TZ3S18	Early calcite	Є2s	5 853.6	-9.6	-2.8
H4S17	Early calcite	Є1x	5 727.4	-8.9	-2.5	0.709 422	±11
TZ1S17	Early calcite	Є3x	6 421.3	-9.4	-2.6	0.709 291	±12
ZS1S15	Early calcite	Є2s	6 521.6	-9.9	-2.3	0.709 112	±10
ZS5S7	Late calcite	Є2a	6 235.4	-11.3	-4.3	0.709 514	±11
Z4S12	Late calcite	Є1w	7 236.5	-11.4	-3.7	0.709 432	±12
H4S16	Late calcite	Є2s	5 448.6	-10.9	-4.4
F1S18	Late calcite	Є1x	4 412.6	-11.1	-3.3	0.709 426	±10
Z4S9	Late calcite	Є2a	6 721.2	-11.7	-3.6	0.709 418	±11
TZ3S10	Late calcite	Є3x	5 436.2	-11.3	-3.3

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Figure  4.  (a) cross-plot of δ¹³C and δ¹⁸O values for limestones, RD1-RD3 and CD dolomite, and early and later-stage calcite cement. The coeval Cambrian seawater isotopic range marked with blue bar (range for seawater δ¹⁸O) and yellow bar (range for seawater δ¹³C) refer to Montañez et al. (2000) and Veizer et al. (1999); (b) cross-plot of ⁸⁷Sr/⁸⁶Sr ratios versus δ¹⁸O values of limestones, RD1-RD3 and calcite cements showing dispersed values with enrichment in ⁸⁷Sr/⁸⁶Sr ratios and decreasing δ¹⁸O values. Note the range for ⁸⁷Sr/⁸⁶Sr values (0.708 872–0.709 400) for Cambrian seawater (e.g., Montañez et al., 2000; Veizer et al., 1999; Burke et al., 1982).

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The RD1 have slightly higher isotopic values than the RD2, RD3 and CD dolomites, with δ¹⁸O values ranging from -9.8‰ to -7.1‰ VPDB (average -8.6‰, n=10) and δ¹³C values from -1.4‰ to -0.4‰ VPDB (average -0.8‰). The ranges of δ¹⁸O and δ¹³C values in RD2, RD3, CD dolomites and early-stage calcite are similar and largely overlap (Table 1), with isotopic values from -10.9‰ to -7.1‰ VPDB (average -9.3‰, n=46) and -3.9‰ to -0.5‰ VPDB (average -3.1‰). In contrast, later-stage calcite has the lowest δ¹⁸O and δ¹³C values compared to those reported above, ranging from -11.7‰ to -10.9‰ VPDB (average -11.3‰, n=6) and δ¹³C values from -4.4‰ to -3.3‰ VPDB (average -3.9‰).

3.2.2   Strontium isotope

The ⁸⁷Sr/⁸⁶Sr ratios for the forty-six analyzed samples range from 0.708 879 to 0.709 514 (Table 1; Fig. 4b) are mostly overlapping with the estimated range of ⁸⁷Sr/⁸⁶Sr ratios for Cambrian seawater (0.708 800–0.709 400; Montañez et al., 2000; Veizer et al., 1999; Burke et al., 1982). The limestone samples yielded a narrow range of ⁸⁷Sr/⁸⁶Sr ratios from 0.708 879–0.709 247 (average 0.709 068, n=5). The RD1-RD3, CD dolomite and early-stage calcite cement have very similar ranges of ⁸⁷Sr/⁸⁶Sr ratios from 0.708 879 to 0.709 422 (n=37). However, the later-stage calcite cement have higher ⁸⁷Sr/⁸⁶Sr ratios compared to those reported above, ranging from 0.709 418 to 0.709 514 (average 0.709 448, n=4).

3.4 Fluid Inclusion Microthermometry

In this study, the fluid inclusion assemblages (FIAs) are defined following the rules of Goldstein and Reynolds (1994). Two phase (liquid+vapor) fluid inclusions were observed along growth and fracture zones in host dolomite crystals, CD dolomite and early- to later-stage calcite with vapor accounting for about 7%–15% of the total inclusion volume. The oil-filled inclusions contain liquid and vapor phases (Fig. 5c). Under fluorescent light, the oil-filled fluid inclusions display blue and bright yellow fluorescence (Fig. 5b) and colorless under transmitted light.

Figure  5.  (a) histogram showing the T_h values of brine in host dolomites, CD dolomite, and early and later-stage calcite cement; (b) photomicrograph of oil inclusions in host dolomite crystal showing bright yellow fluorescence, image taken under fluorescent light; (c) photomicrograph of two-phase brine (red arrows) and oil inclusions (yellow arrows) in CD dolomite cement; (d) cross-plot of T_h values versus salinity values of fluid inclusions showing a nonlinear relationship.

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The ranges of homogenization temperatures (T_h) lies between 94.6 and 164.2 ℃ for CD dolomite, whereas the RD2 and RD3 show a relatively lower T_h values ranging from 80.1 to 141.5 ℃ (Table 2; Fig. 5a). Salinity derived from T_{m ice} for the RD2, RD3 and CD dolomites ranges of 15.4 wt.%–22.6 wt.% NaCleq and 17.7 wt.%–24.5 wt.% NaCleq, respectively. The T_h values for early-stage calcite range from 73.2 to 110.6 ℃, which is slightly lower than that of later-stage calcite with values ranging from 91.4 to 128.7 ℃. Salinity derived from T_{m ice} for early and later-stage calcite overlaps, and lies between 8.6 wt.% and 12.8 wt.% NaCleq (Table 2). The ranges of T_h and salinity of fluid inclusions obtained for this study display a nonlinear relationship (Fig. 5d).

Table  2.  Fluid inclusion data of CD dolomite, early and later-stage calcite cement and host dolomites in the Cambrian carbonate successions, West-central Tarim Basin

Sample code	Mineral Name	FIAs	Number of fluid inclusions	Th (℃)			Tm (℃)			Salinity (wt.% NaCleq)
				Min	Max	Mean	Min	Max	Mean	Min	Max	Mean
Z4S12	Vein (CD)	1	3	144.6	151.2	147.5	-18.6	-21.1	-19.3	22.7	24.8	23.4
Z4S13	Vein (CD)	1	4	146.5	160.8	159.6	-21.7	-19.5	-20.4	19.7	24.6	22.6
T1S6	Vug (CD)	1	4	95.5	113.8	97.4	-22.2	-19.6	-21.3	23.7	26.3	24.5
	Vug (CD)	2	4	94.8	105.8	97.6	-21.3	-18.1	-20.2	19.1	24.7	22.6
TZ1S20	Vein (CD)	1	3	142.1	161.2	151.7	-20.4	-16.7	-18.6	19.6	22.2	20.9
	Vein (CD)	2	5	140.6	151.3	143.2	-15.4	-12.8	-13.6	13.8	23.6	19.8
TZ1S12	Vug (CD)	1	4	95.7	103.5	98.4	-17.4	-15.9	-16.5	17.6	20.7	19.6
	Vug (CD)	2	6	94.8	104.2	96.5	-19.1	-16.4	-17.4	21.6	24.2	22.3
FIS19	Vein (CD)	1	3	123.3	126.8	124.6	-18.6	-16.5	-17.4	17.7	23.6	21.6
	Vein (CD)	2	4	133.4	164.2	152.4	-21.8	-19.1	-20.6	20.3	26.8	22.5
ZS5S6	Vein (CD)	1	3	127.6	141.4	134.4	-18.3	-15.6	-16.6	16.4	23.3	20.4
	Vein (CD)	2	5	121.7	128.5	124.3	-15.3	-12.8	-13.4	15.5	18.8	17.7
ZS1S14	Vein (CD)	1	5	123.6	137.4	131.4	-17.3	-14.6	-15.6	16.5	22.3	20.3
	Vein (CD)	2	4	115.7	125.5	122.4	-18.1	-15.9	-16.5	17.8	20.6	19.7
TZ3S12	Vug (CD)	1	6	118.8	124.2	121.5	-16.3	-14.4	-15.4	17.6	19.2	18.8
	Vug (CD)	2	4	94.6	109.9	97.2	-21.3	-18.1	-20.2	18.6	24.8	22.6
T1S16	Vein (CD)	1	5	142.3	158.8	153.6	-19.4	-17.5	-18.4	19.7	24.6	22.6
	Vein (CD)	2	6	143.4	156.6	152.8	-20.5	-18.6	-19.4	20.8	24.4	21.6
ZS1S10	Vug (CD)	1	4	148.1	155.4	152.7	-22.1	-20.6	-21.2	22.3	25.3	23.2
T1S17	Vug (CD)	1	4	113.2	119.8	115.6	-17.5	-11.8	-15.6	13.6	23.5	19.3
TZ3S18	Early calcite	1	4	86.4	102.8	96.6	-8.6	-7.3	-8.1	9.6	11.4	10.3
	Early calcite	2	3	81.2	96.5	83.3	-11.8	-7.8	-10.1	11.5	13.8	12.8
	Early calcite	3	4	96.7	110.5	101.3	-8.4	-5.8	-6.3	8.7	11.6	9.8
H4S17	Early calcite	1	3	87.2	108.8	96.6	-7.8	-4.8	-5.7	9.3	11.8	10.2
	Early calcite	2	4	93.2	104.6	98.4	-9.2	-5.8	-8.3	9.6	12.7	10.1
	Early calcite	3	4	83.7	101.2	93.6	-10.4	-7.5	-8.3	8.9	11.7	9.8
	Early calcite	4	6	96.5	106.3	97.4	-9.7	-7.6	-6.4	9.6	11.8	10.4
TZ1S8	Early calcite	1	4	89.2	103.1	92.6	-10.1	-8.4	-9.2	10.4	13.8	11.2
	Early calcite	2	3	77.2	97.5	84.6	-7.1	-5.8	-6.4	8.4	10.6	9.3
TZ1S17	Early calcite	1	6	85.4	110.6	91.3	-7.2	-6.1	-6.8	7.8	9.7	8.6
ZS1S15	Early calcite	1	3	87.5	108.4	98.4	-10.7	-5.8	-8.3	8.9	13.7	11.8
	Early calcite	2	4	96.5	105.2	102.4	-10.7	-5.8	-8.3	8.9	14.7	11.8
	Early calcite	3	5	73.2	83.9	77.4	-8.5	-6.3	-7.3	8.6	10.4	9.5
H4S16	Later calcite	1	4	91.4	111.7	98.6	-7.8	-5.8	-5.8	10.4	14.6	11.2
	Later calcite	2	5	94.6	104.8	98.4	-10.4	-7.5	-8.3	9.9	12.7	10.8
	Later calcite	3	4	112.1	122.7	115.4	-10.4	-8.7	-9.5	10.7	13.5	11.8
TZ3S16	Later calcite	1	3	118.6	124.1	121.4	-10.3	-7.4	-8.4	8.7	12.7	9.8
	Later calcite	2	4	101.2	112.5	97.8	-9.1	-6.6	-8.3	11.7	15.6	12.3
	Later calcite	3	4	115.1	125.6	118.3	-10.8	-7.7	-9.4	10.5	12.3	10.2
ZS5S7	Later calcite	1	3	104.6	118.8	111.3	-7.3	-4.7	-6.5	9.4	10.8	9.6
	Later calcite	2	4	114.7	118.4	115.2	-9.5	-7.2	-8.2	10.7	12.9	11.2
Z4S12	Later calcite	1	3	97.6	110.4	106.5	-6.8	-5.5	-6.2	7.5	10.4	9.6
	Later calcite	2	4	110.7	117.4	112.2	-10.2	-7.3	-9.8	10.5	16.4	11.5
Z4S9	Later calcite	1	5	114.7	122.3	117.6	-9.4	-7.5	-8.6	9.8	12.3	10.8
	Later calcite	2	3	117.1	127.6	121.3	-6.8	-3.8	-5.6	8.3	9.8	8.8
F1S18	Later calcite	1	4	121.4	128.7	123.5	-8.6	-6.3	-6.7	8.9	10.6	9.7
	Later calcite	2	5	116.3	122.6	118.5	-9.4	-7.8	-8.6	10.7	14.6	11.2
F1S18	Later calcite	1	4	121.4	128.7	123.5	-8.6	-6.3	-6.7	8.9	10.6	9.7
	Later calcite	2	5	116.3	122.6	118.5	-9.4	-7.8	-8.6	10.7	14.6	11.2
TZ3S15	RD2	1	4	91.3	100.6	94.6	-12.7	-10.9	-11.4	13.7	17.1	15.4
TZ1S20	RD2	1	4	117.9	126.7	123.4	-14.4	-12.5	-13.5	15.6	18.7	17.4
H4S12	RD2	1	5	111.6	125.3	113.2	-17.4	-12.8	-16.6	13.7	23.6	19.2
ZS1S14	RD3	1	4	80.1	96.8	91.6	-16.2	-15.3	-17.4	19.8	23.3	22.5
TZ3S18	RD3	1	5	86.7	117.3	112.4	-15.3	-10.7	-13.6	15.2	19.7	17.3
ZS5S8	RD3	1	4	108.9	122.2	110.3	-14.7	-12.9	-13.6	15.5	20.3	17.4
ZS1S15	RD3	1	6	121.8	131.8	124.7	-17.5	-12.4	-15.5	18.8	20.6	19.7
Z4S12	RD3	1	4	129.2	141.5	134.8	-19.3	-15.6	-17.3	19.4	24.7	22.6

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4. DISCUSSION

4.1 Paragenetic Sequence

The paragenetic sequence of the crystalline dolomites and the cement phases closely associated with fractures and vug-fillings are shown in Fig. 6a. This sequence is recognized based on detailed field observation, core description and microscopic examinations. The diagenetic process can be interpreted to have occurred in three environmental settings, including near-surface (early), shallow-burial (intermediate) and deep-burial (late) environments. Micritization, fibrous, bladed and blocky calcite cementation are early to intermediate stage processes and are defined approximately based on the onset of stylolitization. The deep-burial diagenetic stage is described based on the occurrence of fracture-related CD dolomite and later-stage calcite (Chen et al., 2004). The aforementioned early to intermediate diagenetic characteristics of the studied dolomites can be distinguished mainly in the partially dolomitized samples of RD1 and some samples of RD2, which allowed us to examine the paragenetic sequence of the main diagenetic phases relatively.

Figure  6.  (a) A summarized paragenetic sequence of the Cambrian carbonates in the West-central Tarim Basin, the definition of the various diagenetic stages (e.g., Chen et al., 2004); (b) cross-plot of δ¹⁸O values (VPDB) versus T_h values of dolomites and calcite cement showing the fluid evolution pathways from RD1-RD3, CD dolomite and calcite cement, the contour lines represent δ¹⁸O_water values (VSMOW) in isotopic equilibrium with the dolomites. The dolomite-water equation: 10³ ln α_{dolomite-water}=3.2 (10⁶ T^-2)–3.3 (e.g., Land, 1985) and 10³ ln α_{calcite-water}=2.78 (10⁶ T^-2)–2.89 (e.g., Land, 1983) were applied. The plot shows that the dolomites were formed from ¹⁸O-enriched fluids compared with the estimated isotopic composition of coeval seawater. The inferred isotopic composition of dolomitizing fluid for the dolomites is shown by the dashed rhomb-shapes.

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The occurrence of the RD1 as 'clustered or floating' rhomb-like crystals in matrix along low-amplitude stylolites in grainstone is the earliest crystalline dolomite formed mostly between early and intermediate diagenetic stages (Figs. 3a–3b). The low-amplitude stylolites were probably formed during early to late intermediate burial diagenesis (Chen et al., 2004).

The exact time of the development of RD2 cannot be determined, however, their formation somewhat overlaps with that of RD1. Although they have similar textures, the RD2 have larger crystal sizes and more curved crystal faces, and generally post-date the RD1 (Figs. 3f–3j). This suggests that the RD2 may have formed during late intermediate diagenetic stage at higher-temperature. Figures 3f, 3m and 3n show more tightly packed, curved crystal textures of RD2 and RD3, although very common in the RD3, it confirms that their formation during late intermediate diagenetic stage was at elevated temperatures (Dong et al., 2013b; Chen et al., 2004).

Furthermore, cementation (e.g., dolomite, calcite, gypsum, and pyrite) and bitumen postdate the crystalline dolomites because they partially or completely occlude or lined dissolution pores and fractures within them (Figs. 3a–3n). The presence of pyrite in the dolomites suggests that precipitation is probably from sulphide-rich fluids under reducing diagenetic conditions. The presence of oil inclusions in crystalline dolomites (Fig. 5b) and CD dolomite (Fig. 5c), suggests a concomitant migration of the basinal fluids that precipitated dolomites and CD dolomite with hydrocarbon under higher-temperature conditions (Cai et al., 2008); however, the crystalline dolomites with vugs and fracture-filled bitumen predate the bitumen (Fig. 3e).

Two stages of calcite cements were observed as sets of cross-cutting calcite-filled fractures (Fig. 3n). The early-stage calcite cements are common in dissolved vuggy-pore as fracture-filled calcite cement in the RD1-RD3 (Figs. 3e, 3l, and 3n). The later-stage calcites are mainly fracture-filled cement (Figs. 3j and 3n) and cross-cuts the early-stage calcite cement, indicating that the latter postdate the former (Fig. 3n).

The extensive tectonic and magmatic emplacement during the Early–Late Permian Period should probably be the most suitable scenario for the timing of fracture-related 'geothermal' dolomite formation events in the study area. The fracturing and cross-cutting relationship between the late-stage calcite cement and the fracture-filled CD dolomite across the deeply-buried crystalline dolomites (Figs. 3d, 3e, 3j and 3n) could have been induced by this extensive tectonic and magmatic emplacement (Zhang et al., 2010), which further provided the heat source for the convective-advective basinal fluids flow in the basin. Hence, the deeply-buried crystalline dolomites in the study area predate the tectonics that induced faulting/fracturing and the latest cement infills.

4.2 Origins of Dolomites and Calcite Cements

The main attributes of the different types of dolomite in the Cambrian strata in the West-central Tarim Basin are summarized in Table 3. There are clear differences in the RD1, RD2, RD3 and CD (saddle) dolomite cement, and early and later-stage calcite cements in terms of crystal shape and size, ¹³C, ¹⁸O and ⁸⁷Sr/⁸⁶Sr isotopes, salinities and homogenization temperatures. The nature of the different stages of dolomitization of the Cambrian carbonates in the study area were reconstructed by integrating information on the spatial distribution of sedimentary facies with the diagenetic paragenesis and geochemical results, taking into account the available geologic setting and burial history of the Tarim Basin (Table 3 and Fig. 6a).

Table  3.  Summary of the main attributes of the different types of dolomite and calcite cements in the Cambrian strata in the West-central Tarim Basin

Dolomite type	Occurrence in thin- section	Crystal shape and size	Extinction and CL colour	Homogenization temperature (Th/℃)	Estimated burial depth (m)	δ18O (‰ VPDB)	δ13C (‰ VPDB)	87Sr/86Sr	Salinity	Fluid type	Diagenetic stage(s) and mechanism
RD1	Dolomite crystal	Nonplanara to planars very fine-fine crystalline	Dull red core to bright-red rims	40 to 60 (estimated)	< 1 200	-9.8– -7.1	-1.4– -0.4	0.708 975 to 0.709 331	~3.5 wt.%	Normal seawater	Early-intermediate burial seawater dolomitization, geothermal convection
RD2	Dolomite crystal	Planare (s) fine-medium crystalline	Sharp extinction dull-red core to bright-red rims	94.6 to 123.4	2 000 to 3 000	-10.9– -7.1	-3.9– -0.5	0.708 876 to 0.709 414	~3.5 wt.% and 11.2– 17.4 wt.% NaCleq	Modified seawater and influx of burial brine	Late intermediate burial dolomitization
RD3	Dolomite crystal	Nonplanara to planars, medium-coarse crystalline	Undulose extinction. Purple-red core to bright-red rims	91.6 to 134.8	2 000 to 3 000	-10.4– -7.6	-3.9– -0.4	0.708 992 to 0.709 421	~3.5 wt.% and 17.3–2.6 wt.% NaCleq	Modified seawater and influx of burial brine	Late intermediate burial dolomitization
CD dolomitecement	Cement	Nonplanara, medium-coarse crystalline, cement (or saddle dolomite)	Zonation. Dull-red core to bright-red rims	94.6 to 164.2	> 3 000	-10.6– -7.8	-3.7– -1.5	0.708 964 to 0.709 412	17.7–24.5 wt.% NaCleq	Influx of burial brine	Late intermediate- deep burial dolomitization by geothermal convection
Early-stage calcite cement	Cement	Equant, coarse- crystalline	Non-luminescent	73.2 to 110.6	2 000 to 3 000	-9.9– -8.9	-2.8– -2.3	0.709 112 to 0.709 422	9.3–12.8 wt.% NaCleq	Influx of burial brine	Late intermediate burial stage, geothermal convection
Later-stage calcite cement	Cement	Columnar, coarse- crystalline	Non-luminescent and light orange	91.4 to 128.7	> 3 000	-11.7– -10.9	-4.4– -3.3	0.709 418 to 0.709 514	9.6–12.3 wt.% NaCleq	Influx of burial brine	Late intermediate- deep burial stage, geothermal convection

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4.2.1   Origins of crystalline dolomites

4.2.1.1   Origin of RD1 dolomites

The anhedral-subhedral crystals of RD1 "clustering" in matrix along low-amplitude stratiform stylolite (Figs. 3a, 3b), suggests that their formation may have linked with pressure dissolution. Their development along low-amplitude stylolites was probably the result of compaction due to overburden pressure from overlying sediments and pressure dissolution during early to intermediate burial diagenesis, resulting in stylolitization. Under intense confining and mild compression pressure, Mg ions could have been squeezed out (or expelled) from Mg-rich rocks (e.g., argillaceous limestone, dolomites, shales, etc., Fig. 2a) into remnant Cambrian seawater preserved within the host limestone aquifer, resulting in the precipitation and growth of dolomite crystals along the stylolites (Machel and Buschkuehle, 2008). The clustering and somewhat floating rhomb-like crystals along the stylolite may have been the result of partial dolomitization due to insufficient supply of dolomitizing fluids to the nucleation zones (Sibley and Gregg, 1987). The dolomitizing fluids that precipitated the RD1 during early- to intermediate-burial seawater dolomitization may have been derived from remnant Cambrian seawater preserved in the coeval carbonate coeval carbonate strata (Tucker and Wright, 2009; Machel, 2004). Seawater dolomitization has been reported to occur over a wide range of depths from early to intermediate (shallow) burial, where circulation of seawater induced by geothermal convective fluids flow is considered to be responsible for the formation of massive dolomites (Whitaker and Xiao, 2010; Machel, 2004). In our study, the early-formed dolomites from seawater-dolomitization may have acted as a nucleus for later, more pervasive dolomitization during early intermediate burial diagenesis (Jiang et al., 2014; Machel, 2004). In addition, Whitaker and Xiao (2010) showed through their reactive transport model that dolomite bodies generated by geothermal convection could completely dolomitized large parts of the platform over a relatively prolonged period, for example, 30 Myr. Moreover, Jiang et al. (2014) reported that shallow-buried seawater dolomitization was the main agent in the formation of massive dolomite bodies in the Upper Permian to Lower Triassic dolomite reservoirs in the Sichuan Basin.

The large overlap of isotopic (δ¹⁸O and δ¹³C) values between RD1 and the coeval seawater values (Figs. 4a and 4b; -10.0‰ to -7.0‰ VPDB, e.g., Montañez et al., 2000; Veizer et al., 1999) further confirms that the dolomitizing fluids were presumably derived from modified connate (Cambrian) seawater (Yang et al., 2017). The RD1 have relatively high δ¹⁸O values (-9.8‰ to -7.1‰ VPDB). Assuming the formation temperature of the RD1 to lie between 40–60 ℃ (e.g., Machel, 2004), with a normal geothermal gradient of 30 ℃/km (Zhu et al., 2007) and an estimated burial depth of < 1 200 m in the Middle to Upper Cambrian strata (Fig. 2b; Zhao et al., 2010). Employing the dolomite-water oxygen-isotope fractionation equation of Land (1985), the calculated δ¹⁸O values of the dolomitizing fluid of RD1 was between -6.2‰ and -4.1‰ VSMOW (Fig. 6b), slightly higher than those of the coeval (Cambrian) seawater and corroborates the nature of dolomitizing fluid explain above. Furthermore, their ⁸⁷Sr/⁸⁶Sr ratios generally fall within the range reported for Cambrian seawater (Fig. 4b; Montañez et al., 2000; Veizer et al., 1999), which may have resulted from fluid-rock interaction during progressive burial, providing further constraint on the nature of their dolomitizing fluids.

4.2.1.2   Origin of RD2 and RD3 dolomites

The curved crystal faces and the tightly packed, planar-s crystals texture of RD2 (Figs. 3f–3j) and the nonplanar-a crystals texture of RD3 (Figs. 3m–3p) with irregular overgrowth rims and increasing crystal sizes relative to RD1, indicates that they may have evolved from the RD1 during late intermediate burial setting at higher-temperature either as a result of significant recrystallization upon the earlier formed RD1 or replacement of remaining calcite during late intermediate burial dolomitization (Machel, 2004) by the increasing influx of dolomitizing fluids (Choquette and Hiatt, 2008; Machel, 2004). As mentioned earlier, the burial history of Cambrian strata in the West-central Tarim Basin (Fig. 2b) is consistent with continuous and rapid subsidence after deposition and as a result, the Cambrian carbonate strata reached a depth of about 4 500 to > 6 000 m during the early Ordovician in the Tarim Basin. The highest temperatures that the Cambrian strata reached were from 180 to 240 ℃ (Fig. 2b). Hence, late intermediate burial dolomitization should probably be favoured at such elevated temperature conditions. The T_h defined formation temperatures of RD2-RD3 lies between 80.1 to 141.5 ℃ (Table 2), apparently above the critical roughening temperature (60–95 ℃, e.g., Sibley and Gregg, 1987). However, the RD3 were formed at a higher-temperature relative to the RD2, and their nonplanar-a crystals texture maybe due to rapid disorder during crystal growth (e.g., Chen et al., 2004).

The large overlap of isotopic (δ¹⁸O, δ¹³C and ⁸⁷Sr/⁸⁶Sr) values of RD2 and RD3 (Figs. 4a, 4b) suggests a similar dolomitizing fluid source from which they were formed at different burial temperatures, with the relative preservation of ¹³C-isotopic signatures despite the different replacement processes and later recrystallization. The comparable δ¹³C values of RD2 with those of RD3 may have resulted from biogenic production of CO₂ in soil (Cerling and Hay, 1986), or dissolved carbon derived from the dissolution of host carbonate rocks. On the other hand, the overlapped δ¹⁸O values may have resulted from fluid salinity increase (Fig. 5d), which could have buffered the oxygen isotope fractionation caused by simultaneous rise of fluid temperature (Fig. 6b), declining the δ¹⁸O offset of the RD2 and RD3 dolomites (Figs. 4a and 4b). Moreover, their slightly higher δ¹⁸O values compared with the coeval seawater isotopic range (Fig. 4a) (-10.0‰ to -7.0‰; e.g., Montañez et al., 2000; Veizer et al., 1999), probably resulted from the progressive interaction of basinal fluids with ¹⁸O-rich carbonates and/or silicate host rocks at higher-temperature (Friedman and O'Neil, 1977; Hitchon et al., 1971). Their similar ⁸⁷Sr/⁸⁶Sr ratios presumably resulted from increasing influx of radiogenic strontium from seawater during its interaction with argillaceous limestone and siliciclastic rocks in the Lower Cambrian strata (Fig. 2a). The large overlap of their isotopic values, in addition to their higher T_h and salinity levels (Fig. 5d) suggests that high salinity connate (Cambrian) seawater was the principal dolomitizing fluid (Yang et al., 2017) that precipitated these dolomites during late intermediate burial dolomitization. The widespread occurrence of disseminated pyrite within intercrystalline pores of the RD2 (Fig. 3f), indicates that thermochemical sulphate reduction (TSR) may have taken place in the fluids prior to the precipitation of pyrite crystals at higher-temperatures and under reducing conditions at great depth (Jia et al., 2016; Cai et al., 2004; Van Lith et al., 2003).

The RD2 and RD3 have relatively depleted oxygen isotope values (-10.9‰ to -7.1‰ VPDB). Using the T_h values between 80.1 and 141.5 ℃ for RD2-RD3 and an estimated burial depth of about 2 000 to 3 000 m (Fig. 2b) from the Lower–Upper Cambrian strata, and employing the dolomite-water oxygen-isotope fractionation equation for higher-temperature dolomites (e.g., Land, 1985), the calculated δ¹⁸O values of the dolomitizing fluids for RD2 and RD3 lies between -0.6‰ and 3.8‰ VSMOW (Fig. 6b). Water associated with anhydrite/gypsum dissolution and TSR during late intermediate- to deep-burial dolomitization has been reported to have relatively depleted δ¹⁸O VSMOW values compared to the evolved connate waters (e.g., Jiang et al., 2016, 2015; Worden et al., 1996). The association of hydrocarbon, dissolved anhydrite/ gypsum and H₂S concentration at elevated temperature (> 100 ℃) (Jiang et al., 2016; Cai et al., 2015; Worden et al., 1996), may have generated TSR-rich water or saline connate seawater in the Lower–Middle Cambrian anhydrite/gypsum-rich strata (Fig. 2a) during late intermediate- to deep-burial dolomitization, which could have caused the relatively depleted water-oxygen isotopic valuesin the fluids that precipitated RD2 and RD3 in the study area.

4.2.1.3   Origin of dolomite cement (CD)

The CD (saddle) dolomite predominantly occurs as infills or lining of dissolved vugs (Figs. 3b, 3m, and 3p) and fractures (Figs. 3d and 3l), indicating that their precipitation was closely related to faulting/fracturing at depth. This may be the result of increased pressure gradient in response to enhanced compressional over thrusting that resulted from the advent of the Caledonian orogeny (Gao and Fan, 2014) and thermal anomaly caused by the Early Permian magmatic activity (Yu et al., 2011; Zhang et al., 2010). The higher-temperature expelled basinal brine at depths could have been readily conveyed through the resulting faults/fracture networks as the coeval transient fluids migrated upward. The dissolution of argillaceous limestones, dolomites, mudstone/shale, etc., in the Lower and Middle Cambrian strata (Figs. 2a and 7) by the percolating higher-temperature basinal brine during fluid-rock interaction could have progressively enhanced these basinal fluids with Mg-ion concentrations to become supersaturated with respect to dolomite. As the higher-temperature fluids ascend through faults and related fracture networks, CD (saddle) dolomite could have subsequently precipitated from the higher-temperature Mg-saturated brine-enriched dolomitizing fluid and grow along dissolved vugs and fractures in the dolomite substrates (Davies and Smith, 2006; Al-Aasm, 2003). Merino and Canals (2011) have reported that saddle-shaped dolomite crystals are usually formed owing to increase in Mg²⁺/Ca²⁺ ratio in the dolomitizing fluids, which they attributed the cause to sequestration of Mg ions in dolomite and concomitant dissolution of limestone. This interpretation in the present study is supported by the large overlap of ¹⁸O, δ¹³C, ⁸⁷Sr/⁸⁶Sr, T_h and salinity values of RD2, RD3 and CD dolomites (Figs. 4a, 4b and 5d), in addition to the presence of CD dolomite in vugs and fractures of the crystalline dolomites.

Figure  7.  A conceptual model illustrating fluids flow mechanisms, burial dolomitization as a result of geothermal convection (or Kohout convection, e.g., Kohout et al., 1977); in the Cambrian carbonate successions in the West-central Tarim Basin.

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The large overlap of δ¹⁸O, δ¹³C and ⁸⁷Sr/⁸⁶Sr values of RD2, RD3 and CD dolomites (Figs. 4a–4b), indicates that the higher-temperature brine-enriched dolomitizing fluids are responsible for the precipitation of CD dolomite cement. They were characterized by a similar composition, or presumably inherited the parental signatures of the Mg-rich connate (Cambrian) seawater preserved within the host limestones/dolomites through which it flowed. These fluids were relatively enriched in radiogenic strontium and there probably derived from argillaceous limestones, siliciclastic rocks (Montañez et al., 2000; Veizer et al., 1999). The high salinity levels of CD dolomite (16.2 wt.% to 23.4 wt.% NaCleq; Figs. 5a and 5d), compared to those of the host dolomites, indicates that their dolomitization fluids were more saline than those responsible for the precipitation of the host dolomites. Using the δ¹⁸O values (-10.6‰ to -7.8‰ VPDB), with fluid inclusion-defined temperature range (94.6 to 164.2 ℃) of CD dolomite and employing the fractionation equation of Land (1985), the calculated δ¹⁸O values of the fluids responsible for CD dolomite was between 2.2‰ and 5.8‰ VSMOW (Fig. 6b). This VSMOW values are lower than the calculated water δ¹⁸O derived from seawater (5.5‰ to 9.3‰ VSMOW; e.g., Jiang et al., 2016), and much higher than that of Lower Paleozoic seawater (-11‰ to -8‰ VSMOW; Montañez et al., 2000; Veizer et al., 1999), implying a significant concentration of ¹⁸O in the fluids as a result of fluid-rock interaction with ¹⁸O-rich carbonate and/or silicate minerals within the fault/fracture, and the involvement of saline-rich connate water and possibly TSR-derived water with the dolomitizing fluids that precipitated the dolomite cement at high temperature (e.g., Jiang et al., 2016). In addition, the precipitating fluids would progressively become more saline through water-rock interaction, which is supported by the high T_h values (94.6 to 164.2 ℃) and salinity values (16.2 wt.% to 23.4 wt.% NaCleq) of fluid inclusions in CD dolomite (Fig. 5d). Hence, the δ¹⁸O values of the fluids that precipitated CD dolomite recorded the interactions of both temperature and salinity variations (Fig. 6b) in the deep-buried coeval carbonate aquifer during their upward migration.

According to Davies and Smith (2006), an effective method to evaluate dolomites of hydrothermal origin is to determine when the dolomitization occurred and then compare the burial temperature to the homogenization temperatures of fluid inclusions in the dolomites. Based on the integrated burial history from wells H4 and Z4 (Fig. 2b) for this study, the maximum estimated palaeo-temperatures range from 120 to 240 ℃ for the Cambrian strata. The T_h values of CD dolomite range between 94.6 to 164.2 ℃ mostly lower than the estimated range for palaeo-temperatures. Machel and Lonnee (2002) and White (1957) stated that dolomite should be called hydrothermal only if it can be demonstrated to have formed at a higher-temperature than ambient temperature, regardless of fluid source or drive. This definition does not carry a lower or upper temperature limit. By extension, dolomite cements formed at temperatures lower than ambient are not hydrothermal even if they formed at a rather high temperature (e.g., Machel and Lonnee, 2002). Moreover, dolomite formed in or near thermal equilibrium with the surrounding rocks (e.g., the present study) may be called 'geothermal' dolomites (e.g., Machel and Lonnee, 2002; Packard and Al-Aasm, 2002). Furthermore, Packard and Al-Aasm (2002) reported that the occurrence of saddle dolomites along faults/factures, formed by hot basinal fluids that ascended along fault/fractures from convective fluid flow is 'hydrothermal'. Since most of the CD dolomite in our study was precipitated as cements in the vicinity of vuggy-pores and fractures, they may have been precipitated from the invasion of higher-temperature saline-rich basinal fluids that ascended along faults/fractures networks from convective-advective flow at late intermediate- to deep-burial depths. The formation temperatures of the CD (saddle) dolomite in our study are lower than ambient temperature. Under this circumstance, and based on the aforementioned assertions, we are certain that the growth of CD (saddle) dolomite in the study area are of 'geothermal' origin. This dolomite cement was presumably formed during late intermediate- to deep-burial dolomitization (Fig. 7) (e.g., Machel and Lonnee, 2002) and not necessarily hydrothermal, comparable to reported hydrothermal dolomite occurrence from other places (Guo et al., 2016; Jiang et al., 2015; Dong et al, , 2013a; Davies and Smith, 2006; Chen et al., 2004).

4.2.2   Origin of early and later-stage calcite cement

Two generations of calcite cement infills were recognized in dissolved vugs and fractures (Figs. 3e, 3j, 3l and 3n) in the studied samples. Subsequent to dolomitization and dolomite cementation, the higher-temperature basinal brines presumably underwent progressive cooling with the T_h values of early-stage calcite (73.2–110.6 ℃), later-stage calcite (91.4–128.7 ℃), and a shift towards lower salinity (8.6 wt.%–12.8 wt.% NaCleq) and (8.8 wt.%–12.3 wt.% NaCleq) respectively, and a presumed higher Ca²⁺/Mg²⁺ ratio which resulted in the precipitation of coarse-crystalline equant-columnar calcite-filled vugs and fractures. As earlier mentioned, the increasing Ca²⁺/Mg²⁺ ratios in dolomitizing brines has been attributed to the sequestration of Mg ions in dolomite and the concomitant dissolution of limestone (e.g., Merino and Canals, 2011).

There exists an overlap of isotopic (δ¹⁸O, δ¹³C and ⁸⁷Sr/⁸⁶Sr) values of early-stage calcite cement and CD dolomite; however, an apparent offset exists with those of later-stage calcite cement (Figs. 4a–4b), indicating that the latter was precipitated from higher-temperature basinal fluids enriched in ¹⁸O concentrations, probably after the CD dolomite and early-stage calcite cement. Using the δ¹⁸O values (-10.9‰ to -11.7‰) of later-stage calcite and T_h (91.4–128.7 ℃), and the oxygen isotopes fractionation equation between calcite and water (Land, 1983), it shows that the basinal brines had δ¹⁸O VSMOW values of -2.8‰ to 1.8‰, and the early-stage calcite (δ¹⁸O VSMOW= -3.4‰ to 1.4‰) (Fig. 6b). This range of oxygen isotopic composition inferred from the T_h values of basinal brines can be attributed to various degrees of isotopic evolution by fluid-rock interaction and/or various extents of mixing between meteoric water and evaporative basinal brines. The remarkable decrease in the δ¹⁸O values of later-stage calcite cement compared to early-stage calcite, confirmed their precipitation from diluted (meteoric water influx?) higher-temperature fluid enriched in ¹⁸O concentrations. This interpretation of fluid mixing corroborates the lower salinity levels of the calcite cements relative to CD dolomite and host dolomites, suggesting the incursion of meteoric waters (Zhao et al., 2017) which probably occurred as a result of the progressive tectonic uplift of the Tarim Block during the Early Permian Period (Zhang et al., 2010). Moreover, the high ⁸⁷Sr/⁸⁶Sr ratios of calcite cement suggests an increase influx of radiogenic strontium is probably from leaching of basic magmatic (Sun et al., 2016) or siliciclastic rocks in the study area (Fig. 2a).

4.3 Conceptual Model of Fluid Flow Regimes in the Study Area

The Tarim Basin has experienced multiple stages of tectonic evolutions, including the Caledonian, Hercynian, Indosinian and Himalayan Orogenies (He et al., 2016; Tang, 1997). Under these overall tensional-compressional tectonic regimes, in addition to the intermittent reactivation of basement structures by the regional tectonics during the Late Hercynian orogeny, fault and fracture systems were created in the Lower Paleozoic sedimentary packages (Fig. 7) (Gao and Fan, 2014; Tang, 1997), particularly in the West-central Tarim Basin. The subduction and overthrusting across the convergent plate margin (Fig. 1) caused by the strong tensional-compressional tectonic regimes in the Tarim Basin could have resulted in significant pressure-load and enhancedfluid expulsion by squeezing along the tensional and overthrust fault planes (Fig. 7) (Machel and Cavell, 1999).

In this study, we considered two main regimes of fluid flow in the deeply buried Cambrian carbonate successions, which were presumably linked to the tectonic evolution of the sedimentary basin. The first regime involved influx of magnesium-rich higher-temperature basinal brines (T_h: 80.1–164.2 ℃; salinity 15.4 wt.%–24.5 wt.% NaCleq) into the Cambrian successions along deep-seated faults/fractures. This regime was probably active during the development of the strong tensional-compressional tectonic regimes that resulted from the advent of the Caledonian orogeny (Upper Cambrian to Lower–Late Ordovician) and Hercynian orogeny (Early–Late Permian) (He et al., 2016; Gao and Fan, 2014; Tang, 1997), causing the formation of tensional and thrust faults, as well as cementation by CD dolomite. The tensional and thrust faults, fracture networks and permeable horizons (bedding planes, stylolites, and unconformity surfaces) presumably provided conduits for the cross-formational and ascending higher-temperature magnesium-rich basinal brines from deep over-pressured zones to shallower depths resulting in the precipitation of CD (saddle) dolomite from convective-advective flow into the fractures and vuys of the host carbonates. The large overlap of δ¹⁸O, δ¹³C, high T_h and salinity values of CD dolomite and host dolomites, supports the assertion that the CD dolomite was precipitated from higher-temperature basinal fluids diluted with high-salinity brines, since the fractionation of oxygen isotopes between dolomite and fluids is temperature dependent (Mansurbeg et al., 2016).

The second fluid flow regime in the study area involved the downward incursion of meteoric waters, mixing with ascending higher-temperature basinal brines during the tectonic uplift events in the basin, probably resulting in the progressive cooling of the basinal brines (T_h =73.2–128.7 ℃; early- to later-stage calcite cement) and a shift towards lower salinity (8.6 wt.%–12.8 wt.% NaCleq). Subsequent to dolomite cementation (Mg²⁺ consumption?), which was concomitant with hydrocarbon migration along vugs and fracture networks, resulted in increase in Ca²⁺/Mg²⁺ ratio in the fluids (Merino and Canals, 2011), and may have promoted the precipitation of calcite cement in vugs and fractures possibly before or during the advent of the Late Hercynian orogeny.

Although there exists an overlap of ⁸⁷Sr/⁸⁶Sr ratios between the RD1-RD3, CD dolomites and early-stage calcite cement, an apparent offset occurs with those of later-stage calcite cement (Fig. 4b). In addition, the later-stage calcite cement in this study has the highest ⁸⁷Sr/⁸⁶Sr ratios (0.709 418–0.709 514; Fig. 4b), suggesting input of ⁸⁷Sr from the interaction of higher-temperature brine-enriched basinal fluids with basic magmatic and siliciclastic rocks particularly in the Cambrian strata, further constraining the source of fluids flow in the study area.

5. CONCLUSIONS

The following conclusions were drawn from petrographic, geochemical (O-C-Sr isotopes) and fluid inclusion results of the deeply buried Cambrian carbonate successions in the West-central Tarim Basin.

1. Three types of crystalline dolomites (RD1–RD3), CD dolomite, and early and later-stage calcite cement are distinguished. The occurrence of RD1 along low-amplitude stylolite suggests that their formation may have linked with pressure dissolution by which minor Mg ions were probably released for replacive dolomitization during early- to intermediate-burial seawater dolomitization. The tightly packed and curved crystal faces of RD2 and RD3 with irregular overgrowth rims and increasing crystal sizes relative to RD1 may have resulted from intense recrystallization and replacement upon the RD1 or from the remaining precursor limestones by the influx of dolomitizing fluids at higher-temperatures during late intermediate burial dolomitization.

2. The large overlap of δ¹⁸O, δ¹³C and⁸⁷Sr/⁸⁶Sr values of RD1, RD2, RD3 and CD dolomite with coeval seawater values, suggests that the principal dolomitizing fluids that precipitated these dolomites was connate (Cambrian) seawater preserved in the host limestones/dolomites. Their ⁸⁷Sr/⁸⁶Sr ratios suggest fluid-rock interaction and influx of radiogenic strontium into the connate seawater.

3. Two main regimes of fluid flow into the deeply buried Cambrian carbonate succession are considered, which are linked to the tectonic evolution of the basin. The first regime involved influx of magnesium-rich higher-temperature basinal brines into the Cambrian carbonate successions along deep-seated faults/fractures, resulting in cementation by CD dolomite, presumably during the advent of the Caledonian and Hercynian orogenies. The second regime involved the incursion of meteoric waters, mixing with ascending higher-temperature basinal brines during the tectonic uplift event in the Tarim Basin, probably resulting in the progressive cooling of the basinal brine and a shift towards lower salinity. Subsequently, an increased in Ca²⁺/Mg²⁺ ratio in the fluids could have promoted the precipitation of calcite cement in vugs and fractures from geothermal convective-advective flow at late intermediate-to deep-burial depths in the study area.