Genesis and Reservoir Implications of Multi-Stage Siliceous Rocks in the Middle–Lower Ordovician, Northwestern Tarim Basin

Abstract

Siliceous rocks of various colors and types are extensively developed within the Middle–Lower Ordovician carbonate along the Northwest Tarim Basin. Their genesis provides important insights into the evolution of basinal fluids and the associated diagenetic alterations of the carbonates. Based on petrographic, geochemical, fluid inclusion, and petrophysical analyses, this study investigates the origin of siliceous rocks within the Middle–Lower Ordovician carbonate formations (Penglaiba, Yingshan, and Dawangou formations) in the Kalpin area, Tarim Basin, and investigates the impact on hydrothermal reservoirs. The results reveal two distinct episodes of siliceous diagenetic fluids: The first during the Late Ordovician involved mixed hydrothermal fluids derived from deep magmatic–metamorphic sources, formation brines, and seawater. Characterized by high temperature and moderate salinity, it generated black chert dominated by cryptocrystalline to microcrystalline quartz through replacement processes. The second episode developed in the Middle–Late Devonian as a mixture of silicon-rich fluids from deep heat sources and basinal brines. In conditions of low temperature and high salinity, it generated gray-white siliceous rocks composed of micro- to fine crystalline quartz, spherulitic-fibrous chalcedony, and quartz cements via a combination of hydrothermal replacement and precipitation. A reservoir analysis reveals that the multi-layered black siliceous rocks possess significant reservoir potential amplified by the syndiagenetic tectonic fracturing. In contrast, the white siliceous rocks, despite superior petrophysical properties, are limited in scale as they predominantly infill late-stage fractures and vugs, mainly enhancing local flow conduits. Hydrothermal alteration in black siliceous rocks is more intense in dolostone host rocks than in limestone. Thus, thick (10–20 m), continuous black siliceous layers in dolostone and the surrounding medium-crystalline dolostone alteration zones, are promising exploration targets. This study elucidates the origins of Ordovician siliceous rocks and their implications for carbonate reservoir properties. The findings may offer valuable clues for deciphering the evolution and predicting the distribution of hydrothermal reservoirs, both within the basin and in other analogous regions worldwide.

1. Introduction

Siliceous rocks, as a distinct type of sedimentary rock, are globally distributed in marine strata across various geological periods. Their formation is closely associated with specific paleo–oceanic environments [1,2], biological activity [3], or hydrothermal processes [2,4,5]. These rocks not only serve as important indicators for reconstructing paleo–oceanic conditions [5] and plate tectonic activities [6], but also hold significant petroleum geological potential. Biogenic siliceous rocks can be high-quality source rocks and reservoirs, whereas hydrothermal siliceous rocks and associated succession often exhibit reservoir–seal potential [7,8]. Recent discoveries of large siliceous rock hydrocarbon reservoirs in North America, such as Wolf Springs Field [9], Bedford Devonian Field [10], and Parkland Field [11], confirm their exploration value.

The origins of silica in carbonate rocks are diverse, each with distinct silica sources and resulting rock characteristics. Although methods exist to identify genetic types of siliceous rocks using petrological, elemental geochemical [11,12], and isotopic geochemical criteria [13], different indicators often yield conflicting interpretations. Furthermore, siliceous rocks of different genetic types have varying implications for hydrocarbon accumulation in carbonate rocks. However, in practical exploration, most studies still evaluate siliceous rocks in carbonate succession as products of a single silicification event, lacking systematic distinction of the reservoir effects of different genetic types.

Continuous breakthroughs in deep to ultra-deep petroleum exploration in the Tarim Basin have revealed abundant siliceous rocks in the Middle–Lower Ordovician carbonate rocks across most basin areas (e.g., Shunnan, Gucheng, Tahe, Shuntuoguole, and Fuman) [14,15,16,17]. Frequent industrial gas flows from siliceous rocks and adjacent succession indicate promising exploration prospects, attracting growing scholarly attention. Most researchers currently attribute these siliceous rocks to having a hydrothermal origin [14,15,16,17,18,19], yet several controversies persist. First, debate continues on whether they formed through hydrothermal replacement or sedimentation [18,19,20]. Second, ambiguous deep fluid sources and diagenetic timing lead to ambiguity in the formation age of Ordovician siliceous rocks. While some studies support pre-Permian formation [14,17], others associate it with the Permian magmatic event [15,18,21]. Third, poorly constrained timing and triggers of hydrothermal activity hinder systematic research on the self-reservoir properties of siliceous rocks and fluid alteration effects on surrounding rocks, limiting prediction accuracy for hydrothermal reservoirs.

The northwestern margin of the Tarim Basin exhibits a wide range of siliceous rock types within its Ordovician outcrops [22,23,24]. This study samples Middle–Lower Ordovician siliceous rocks and their host rocks from outcrop sections on the northwestern margin of the Tarim Basin and conducts comprehensive research, including petrographic, geochemical, fluid inclusion microthermometric, and reservoir characteristic analyses. The objective of this study is to reveal the origin and diagenetic evolution of siliceous rocks, thereby evaluating the impact of hydrothermal siliceous activity on carbonate reservoir formation, and to provide a methodological framework for investigating the genesis and reservoir effects of siliceous rocks in other basins.

2. Geological Settings

The Kalpin Fault Uplift is situated in the basin–range junction zone between the Paleo–Tianshan Mountains and the northern Tarim Plate in the northwestern Tarim Basin [25]. As a first-order tectonic unit, it closely connects with the Wushi and Kuqa depressions and covers about 20,000 km² (Figure 1a). The area is divided into three regions from west to east: Xike’er, Kalpin, and Aksu, and features multiple NE-trending major faults with a structural framework of stacked thrust sheets. The regional tectonic system is complex and can be divided, based on structural zoning, into the Wensun Uplift, Shajingzi Structural Belt, and Kalpin Thrust Belt. The Kalpin Thrust Belt is dominated by imbricate thrust nappe structures, forming typical compressional geomorphological features such as tectonic windows and klippes [26,27] (Figure 1b). The region has undergone multiple tectonic phases. Early Caledonian–Hercynian N–S compression formed basement thrust faults, while the Himalayan India–Eurasia collision later intensified thrusting, nappe emplacement, strike-slip motion, and hydrothermal alteration [27,28,29,30]. Succession from the Upper Cambrian to Permian is exposed, with complete Middle–Lower Ordovician boundaries of the Penglaiba, Yingshan, Dawangou, and Saergan formations visible in classic Ordovician sections [31]. The characteristics of siliceous rocks within the carbonate are exceptionally clear [20,22] (Figure 1c). The Cement Plant Section (Kamatikan section) [32] lies within the Yimugantawu Fault Zone, near the Subashi Reservoir, about 15 km northwest of Kalpin County, and serves as a key Ordovician outcrop for geological research in the Tarim Basin [33]. Sedimentary facies show the transition from the Upper Cambrian Xiaqiulitag Formation platform margin to a Middle–Lower Ordovician shallow marine open platform. The Penglaiba Formation consists mainly of limestone, with black dolomite in the middle portion; the Yingshan Formation contains sparry calcarenite and micritic limestone; and the Dawangou Formation is made up of nodular limestone [34,35]. Multiple phases of hydrothermal activity have overprinted the carbonate rocks, producing unevenly distributed layered, banded, and nodular siliceous rocks that enhance reservoir heterogeneity and pore connectivity [36].

3. Samples and Methods

3.1. Samples

We systematically collected siliceous rocks and their surrounding carbonates from the Middle–Lower Ordovician Penglaiba, Yingshan, and Dawangou Fm in the Kalpin Cement Plant Section, Kalpin Fault Uplift, northwestern Tarim Basin. A total of 15 siliceous rock and 120 carbonate rock samples were collected. Siliceous rocks with diverse morphologies occur mainly in Layers 14, 24, 37, and 58 of the Penglaiba Fm; the middle part of the Yingshan Fm; and the base of the Dawangou Fm (Figure 1c). Sampling was conducted on the siliceous rock-bearing intervals. Carbonate samples were taken above and below these intervals until reaching zones free of siliceous rocks and hydrothermal alteration. Carbonates were thin-sectioned and microscopically examined to verify the absence of silicifying hydrothermal effects. The spacing between carbonate rock samples was 5–10 cm. Siliceous rocks include black and white varieties—layered, banded and nodular—all sampled for this study.

3.2. Experiment

This study begins with the preparation and petrographic identification of thin sections from siliceous rocks of varying colors and types, aiming to characterize their lithological features. Representative samples—exhibiting distinct colors and occurring in stratified, banded, and nodular forms—were then selected from typical intervals of the Penglaiba, Yingshan, and Dawangou Fm. These samples were subjected to a series of analytical techniques, including major and trace element analysis, rare earth element (REE) analysis, silicon–oxygen isotope analysis, and fluid inclusion microthermometry. Subsequently, reservoir characteristics of the host rocks adjacent to different types of siliceous rocks across various stratigraphic levels were examined, accompanied by porosity and permeability measurements, to evaluate the impact of siliceous rocks on reservoir modification.

Blue epoxy-impregnated thin sections were prepared from 85 carbonate rock samples. Petrography was performed using an Olympus BX53 microscope (Olympus Corporation, Tokyo, Japan) at 22 °C, with pores identified by their blue color under polarized light. A total of 15 siliceous rock and 30 carbonate rock samples were argon-ion polished for microstructural analysis using an FEI Quanta 650 FEG SEM (FEI Company, Hillsboro, OR, USA). Microstructural reservoir characteristics were observed using an FEI Quanta 650 FEG SEM. Cathodoluminescence (CL) observation was conducted on 15 carbonate thin sections with a CL8200 MK5 instrument (Cambridge Image Technology, Hertfordshire, UK), operating at a working current of 250 μA at 11 kV. All these analyses were completed at Southwest Petroleum University. Analysis of 20 samples for major, trace, and rare earth elements, along with Si and O isotopes from 10 siliceous rocks, were analyzed at Chengdu Pupu Testing Technology Co., Ltd. (Chengdu, China). Approximately 1 g of sample powder was digested in a PTFE beaker with HNO₃-HF-HClO₄ (5:5:1) at 250 °C. Subsequently, 8 mL of aqua regia (1:1) was added and heated until the solution volume was reduced to 2–3 mL. The inner walls of the beaker were rinsed with approximately 10 mL of deionized water, followed by gentle heating for 5 to 10 min until the solution became clear, before removal from the hotplate. After cooling, the solution was transferred to a 10 mL graduated polyethylene test tube with a stopper, diluted to the mark with deionized water, thoroughly shaken, and allowed to settle. A 1 mL aliquot of the clear supernatant was pipetted into another polyethylene tube, diluted to 10 mL with 3% (v/v) nitric acid, mixed thoroughly, and then analyzed. Major elements were analyzed using Inductively Coupled Plasma Optical Emission Spectrometry (ICP-OES) using a PerkinElmer 8300V instrument (PERKINELMER, Waltham, MA, USA). Trace elements were analyzed with Inductively Coupled Plasma Mass Spectrometry (ICP-MS) using an Agilent 7700 instrument (Agilent Technologies, Santa Clara, CA, USA). Each element was measured three times. Rare earth element (REE) data were normalized to Post-Archean Australian Shale (PAAS). Ce and Eu anomalies were calculated as (Ce/Ce*)_N = Ce_N/(0.5La_N + 0.5Pr_N) and (Eu/Eu*)_N = Eu_N/(Sm_N × Gd_N)^1/2 [37].

For silicon isotope analysis, 5 mg of sample powder was uniformly mixed with 200 mg of NaOH powder, fused at 730 °C for 10 min, and cooled. The resulting material underwent water extraction and acid dissolution, with the solution pH adjusted to 2.2 for purification. Silicon isotope purification was performed using AG 50W-X12 (200–400 mesh) cation exchange resin. The resin was sequentially cleaned with various acids and equilibrated with water before sample loading. The silicon fraction was ultimately collected using Milli-Q (MQ) water. Measurements were then conducted using a Neptune Plus multi-collector inductively coupled plasma mass spectrometer (MC-ICP-MS) (Thermo Fisher Scientific, Waltham, MA, USA). The international reference material GBW04421 (quartz) was used as the monitoring standard during analysis, with an analytical error within ±0.1‰. For oxygen isotope analysis, the BrF₅ fluorination method was employed to determine the oxygen isotope composition of oxide and whole-rock samples. After specific pretreatment (baking, acid washing to remove carbonates and/or organic matter), samples were reacted with BrF₅ in a vacuum system at different temperatures (550 °C for 6 h for low-temperature minerals, 630 °C for 14 h for high-temperature minerals, 580 °C for 8–10 h for whole rocks). The liberated oxygen was cryogenically separated using liquid nitrogen and collected. Measurements were performed using a Thermo Fisher Scientific 253 Plus gas isotope ratio mass spectrometer (Thermo Fisher Scientific, Waltham, MA, USA), relative to the V-SMOW standard. The analytical precision (1σ) for a single sample was ±0.05‰, and GBW04421 was used as the working standard for quality control.

Systematic fluid inclusion analysis was completed at the Beijing Research Institute of Uranium Geology. Inclusion types were first identified and marked on the thin sections using a ZEISS Imager.M2m microscope (Carl Zeiss AG, Oberkochen, Germany). The marked sections were then placed in a LINKAM THMS 600 heating–freezing stage (Linkam Scientific Instruments Ltd., Surrey, UK) (precision 0.1 °C). The temperature within the heating–freezing stage was gradually increased until the two-phase aqueous inclusions reached complete homogenization, and the homogenization temperature (Th) was recorded. Subsequently, the stage was cooled to −120 °C to completely freeze the inclusions, followed by gradual rewarming. The temperature at which the last ice crystal disappeared was recorded as the final ice melting temperature (Tm). Salinities were calculated from the Tm values using the Bodnar [38] equation. Physical property analysis of the siliceous rocks and surrounding carbonate reservoirs was performed by the National Experimental Teaching Center for Oil and Gas Geology and Exploration at Southwest Petroleum University. Analyses were conducted using a CMS-300 unconventional porosity and permeability analyzer (Core Lab, Houston, TX, USA). Based on Boyle’s law, the porosity and permeability are calculated from the pressure change and the gas volume when a known quantity of standard gas, under a set initial pressure, undergoes isothermal expansion into a core chamber initially at atmospheric pressure, allowing the gas to diffuse into the pore spaces of the rock sample.

4. Results

4.1. Petrography

Within the Middle–Lower Ordovician carbonate rocks of the northwestern Tarim Basin (Penglaiba to Dawangou Fm), siliceous rocks occur in layered, banded, and nodular forms. The gray-black variety exhibits all these morphologies, while the white variety is present only as banded and nodular types. The lateral and vertical distribution of black siliceous rocks is significantly more extensive. Petrographically, the black rocks are predominantly cryptocrystalline to microcrystalline chert, whereas the gray-white rocks consist mainly of microcrystalline quartz, chalcedony, and granular quartz cement.

4.1.1. Lithofacies

(1)

Penglaiba Formation (O₁p)

The siliceous rocks in the Penglaiba Fm are primarily black and gray-white in color, with some black varieties altered to brownish-brown by late-stage oxidation (Figure 2a). In the lower part of the formation, black siliceous rocks occur as thin-bedded and stratiform-banded types along dolomite bedding planes and fracture zone peripheries. Upwards, they transition into discontinuous bands and nodules concentrated in fracture clusters. Thin-bedded black siliceous rocks mainly develop at lithological interfaces between fine-crystalline dolomite and sparry calcarenite–dolomicrite where bedding-plane fractures are relatively wide. These layers are 5–10 cm thick, locally reaching 15–20 cm, and exhibit sharp, often undulating contacts with the surrounding carbonate country rock (Figure 2b). Banded black siliceous rocks develop along dolostone bedding planes and fracture zones. Individual layers between beds are 5–10 cm thick and 1–2 m long (Figure 2b). In fractured dolostone intervals, bands are irregular, less than 10 cm thick, and laterally discontinuous, with lengths of 0.3–1.5 m (up to 1–2 m; Figure 2b,c). Black nodular siliceous rocks are commonly embedded within carbonate rocks containing well-developed bedding-parallel or tectonic fractures, with dimensions of 5–10 cm × 3–10 cm (Figure 2d), and locally exhibit dissolution-enlarged features.

On an outcrop scale, white siliceous rocks are subordinate respective to black types and are scarce in the Yingshan and Dawangou Fm. They occur mainly as nodules and irregular bands within fractures and dissolution cavities in carbonates or along fractures between black siliceous rocks and the hosting, surrounding rock (Figure 2e,f). Bands are approximately 0.4–1.5 m long and 5–15 cm thick; nodules measure 5–10 cm × 3–5 cm. Additionally, minor amounts of columnar quartz grains with greasy luster are observed in carbonate fractures and dissolution cavities, with crystal sizes of about 1 cm (Figure 2g).

(2)

Yingshan Fm (O_1–2y)

Siliceous rocks in the Yingshan Fm occur predominantly as black bands and nodules, developing within sparry calcarenite and limestone fracture zones (Figure 2h). Their size is smaller than in the Penglaiba Fm. Compared with the Penglaiba Fm, the siliceous rocks in the Yingshan Fm exhibit numerous but more discontinuous bands, with lengths of approximately 0.5–2 m and thicknesses of 3–5 cm. Black siliceous nodules are also more developed, often distributed stratiformly along fractures with sizes of 1–10 cm × 1–10 cm.

(3)

Dawangou Fm (O₂d)

Siliceous rocks in the Dawangou Fm occur as thin black layers, stratiform bands, and nodules within the sparry calcarenite interval. The black thin-layered siliceous rocks are concentrated at the base, comprising approximately four to five layers, each with a thickness of 2–5 cm and relatively extensive lateral continuity. Upwards, they pass to black bands or black nodules (Figure 2i).

4.1.2. Microscopic Characteristics

(1)

Black siliceous rocks

Black siliceous rocks exhibit a high degree of replacement. They contain relict dolomite and bioclasts replaced by quartz (Figure 3a), as well as ghost structures of carbonate intraclasts (Figure 3b). Some uniformly replaced layers preserve the original laminated carbonate texture (Figure 3c). Quartz is predominantly cryptocrystalline (CQ) and microcrystalline (Q1, diameter 0.5–30 μm), with a total content > 95%. These quartz phases are euhedral to anhedral and show wavy extinction (Figure 3a,b). Chalcedony, mainly spherulitic (SC) exhibiting wavy extinction, fills fractures and dissolution pores (Figure 3d).

(2)

White siliceous rocks

White siliceous rocks consist mainly of microcrystalline quartz, chalcedony, and quartz cement. This three-stage sequence fills fractures and vugs (Figure 3e,f). First-generation microcrystalline quartz (Q2, 2–20 μm) forms at fracture edges and in the matrix. Second-generation chalcedony includes spherulitic (SC) and long-fibrous (LC) types (Figure 3f,g). Spherulitic chalcedony (SC) exhibits undulose extinction due to its composite fan-shaped crystal structure, making individual crystal boundaries difficult to distinguish. The crystal sizes range from 10 to 100 μm, and fill the fractures (Figure 3e,f) or dissolution pores. Long-fibrous chalcedony (LC) develops radially from the carbonate matrix at fracture edges towards the fracture center with lengths of approximately 200–700 μm and shows wavy extinction under cross-polarized light (Figure 3f,g). The quartz grain size gradually increases towards the centers of dissolution pores and fractures. Agate-like banding with dolomite (layer thickness < 1 mm) occurs in some samples (Figure 3h), indicating periodic precipitation from siliceous and Mg-rich fluids. Third-generation highly euhedral phanerocrystalline quartz (Q3, 75–100 μm) fills the central parts of pores and fractures (Figure 3i).

4.2. Geochemical Characteristics

4.2.1. Characteristics of Major and Trace Element

The major and trace element compositions of the siliceous rocks are listed in

Table 1. Major element contents and parameters of the Middle–Lower Ordovician siliceous rocks from the Cement Plant Section, Kalpin area.

Sample	Depth (m)	Type	SiO2	Al2O3	Fe2O3	MgO	CaO	Na2O	K2O	MnO	TiO2	P2O5	Al/(Al + Mn + Fe)
			%	%	%	%	%	%	%	%	%	%
O1p-1-n-B	489	Black nodule	85.86	0.24	0.55	0.18	6.94	0.17	0.07	0.04	0.03	0.00	0.23
O1p-14-l-B	443	Black layer	94.46	0.31	1.16	0.44	1.89	0.18	0.06	0.04	0.02	0.00	0.16
O1p-24-l-B	420	Black layer	94.57	0.42	1.99	0.10	0.89	0.11	0.06	0.15	0.03	0.02	0.13
O1p-24-b-B	420	Black band	89.13	0.17	1.07	1.51	3.81	0.11	0.04	0.04	0.02	0.01	0.10
O1p-37-b-B	355	Black band	87.66	0.27	1.02	1.10	4.97	0.15	0.07	0.04	0.03	0.00	0.16
O1p-58-n-B	315	Black nodule	95.63	0.33	1.66	0.11	0.46	0.09	0.05	0.12	0.02	0.02	0.12
O1–2y-b-B	140	Black band	94.58	0.21	2.53	0.04	0.78	0.09	0.04	0.19	0.02	0.02	0.05
O1–2y-n-B	140	Black nodule	84.69	0.24	1.90	0.18	6.18	0.10	0.04	0.14	0.02	0.02	0.08
O2d-l-B	18	Black layer	93.21	0.26	2.20	0.04	1.90	0.09	0.05	0.16	0.02	0.02	0.08
O2d-n-B	17	Black nodule	95.90	0.23	1.14	0.17	1.51	0.14	0.05	0.04	0.03	0.01	0.13
O1p-14-b-W	441	White band	84.85	0.17	0.17	0.15	7.79	0.11	0.03	0	0.02	0.02	0.42
O1p-24-n-W	420	White nodule	94.32	0.20	0.71	0.71	1.48	0.07	0.04	0.05	0.02	0.02	0.17
O1p-55-b-W	333	White band	97.01	0.22	1.19	0.09	0.31	0.06	0.05	0.09	0.02	0.02	0.11
O1p-58-n-W	323	White nodule	93.32	0.34	1.47	0.87	1.40	0.06	0.07	0.10	0.03	0.02	0.14
O1p-58-Q-W	315	Quartz grains	74.04	0.20	0.86	5.58	7.30	0.05	0.06	0.07	0.02	0.02	0.14

and

Table 2. Trace element contents and geochemical parameters of the Middle–Lower Ordovician siliceous rocks from the Cement Plant Section, Kalpin area.

Sample	Depth (m)	Type	Li	Be	Sc	V	Cr	Co	Ni	Cu	Zn	Ga	Rb	Sr	Y	Zr	Mo	Cs	Ba	Th	U	Th/U
			μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g
O1p-1-n-B	489	Black nodule	10.26	0.14	0.22	2.14	4.50	0.73	2.53	5.26	13.15	0.53	1.34	58.65	0.73	3.39	0.38	0.41	0.12	27.36	0.40	0.87
O1p-14-l-B	443	Black layer	4.35	0.06	0.15	1.92	4.38	0.59	2.61	5.34	24.23	0.52	1.14	22.00	0.30	2.68	0.29	2.20	0.12	38.86	0.19	0.08
O1p-24-l-B	420	Black layer	6.90	0.14	0.31	6.23	9.66	1.92	4.11	6.32	17.64	0.80	1.49	23.68	0.62	10.66	0.49	3.99	0.14	42.90	0.47	0.28
O1p-24-b-B	420	Black band	5.68	0.03	0.18	3.76	4.89	0.54	2.04	4.35	16.44	0.53	1.08	30.78	0.41	2.81	0.33	1.04	0.10	13.13	0.28	0.39
O1p-37-b-B	355	Black band	4.27	0.12	0.21	3.32	6.79	1.05	3.59	6.22	23.69	0.56	1.50	24.61	0.60	3.93	0.33	2.34	0.17	13.47	0.31	0.08
O1p-58-n-B	315	Black nodule	3.05	0.12	0.20	5.42	5.54	1.55	3.81	5.88	8.41	0.64	1.38	42.62	0.48	4.69	0.40	3.86	0.19	12.41	0.33	0.07
O1–2y-b-B	140	Black band	6.71	0.02	0.11	6.51	11.20	2.01	4.59	7.55	10.78	0.88	1.08	32.88	0.43	7.44	0.43	1.46	0.26	314.53	0.23	0.54
O1–2y-n-B	140	Black nodule	8.78	0.06	0.20	5.33	7.09	1.81	4.01	6.51	8.20	0.57	1.26	59.72	0.90	4.39	0.37	1.30	0.14	85.63	0.23	0.54
O2d-l-B	18	Black layer	3.98	0.10	0.20	5.26	8.20	1.99	4.27	7.78	7.30	0.69	1.32	62.91	0.61	4.79	0.43	1.44	0.15	45.65	0.30	0.49
O2d-n-B	17	Black noule	4.20	0.05	0.16	4.34	5.29	0.74	3.18	5.49	20.25	0.39	1.23	27.75	0.32	3.64	0.35	2.83	0.17	28.07	0.27	0.08
O1p-14-b-W	441	White band	6.73	0.05	0.16	4.50	1.48	0.20	0.86	0.43	5.26	0.22	0.95	68.82	0.31	2.95	0.19	0.40	0.12	8.18	0.12	0.40
O1p-24-n-W	420	White nodule	1.74	0.09	0.17	6.07	3.18	0.78	2.47	3.45	5.35	0.59	0.88	17.83	0.32	4.44	0.32	2.91	0.09	7.96	0.24	0.12
O1p-55-b-W	333	White band	1.12	0.12	0.21	5.35	4.67	1.06	2.69	4.68	6.64	0.62	0.98	16.29	0.38	4.79	0.35	1.32	0.09	13.15	0.29	0.11
O1p-58-n-W	323	White nodule	2.81	0.11	0.29	5.75	5.14	1.40	3.22	5.60	10.37	0.69	1.58	16.23	0.32	5.34	0.44	1.52	0.11	15.05	0.30	0.12
O1p-58-Q-W	315	Quartz grains	1.84	0.02	0.24	16.27	7.68	0.88	2.14	3.69	5.87	0.46	1.22	25.69	0.41	4.55	0.46	1.01	0.07	9.84	0.32	0.19

. The SiO₂ content ranges from 74.04% to 97.01%, with black varieties ranging from 84.69% to 95.90% and white varieties from 74.04% to 97.01%. In black siliceous rocks, Fe₂O₃ is the dominant component other than SiO₂, ranging from 0.55% to 2.53% (average: 1.52%). In white siliceous rocks, CaO is highest among the other components, ranging from 0.31% to 7.79% (average: 3.66%). MnO content averages 0.10% (0.04%–0.19%) in black rocks and 0.06% (0.001%–0.10%) in white rocks. Both rock types show similarly low Al₂O₃ (0.17%–0.42%) and TiO₂ (0.02%–0.03%). The Al/(Al + Fe + Mn) ratio is 0.05–0.23 in black siliceous rocks and 0.11–0.42 in white siliceous rocks.

Among trace elements, black siliceous rocks contain 4.38–11.20 μg/g Cr (avg. 6.75 μg/g) and 2.04–4.59 μg/g Ni (avg. 3.47 μg/g) and show Th/U ratios of 0.07–0.87. White siliceous rocks contain 1.48–7.68 μg/g Cr (avg. 4.43 μg/g), 0.86–3.22 μg/g Ni (avg. 2.28 μg/g), and Th/U ratios of 0.11–0.40.

4.2.2. Characteristics of Rare Earth Element

The analytical results of rare earth elements for 15 siliceous rock samples from the Penglaiba Fm to Dawangou Fm are shown in

Table 3. Rare earth element (REE) contents and geochemical parameters of the Middle–Lower Ordovician siliceous rocks from the Cement Plant Section, Kalpin area.

Sample	Depth (m)	Type	La	Ce	Pr	Nd	Sm	Eu	Gd	Tb	Dy	Ho	Er	Tm	Yb	Lu	ΣREE	LREE/ HREE	Y/Ho	La/Ho	δCe	δEu	(La/Yb)N
			μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g	μg/g
O1p-1-n-B	489	Black nodule	0.808	1.290	0.165	0.664	0.130	0.034	0.140	0.023	0.115	0.027	0.066	0.011	0.065	0.011	3.55	6.75	26.99	29.72	0.81	1.18	0.91
O1p-14-l-B	443	Black layer	0.488	0.901	0.101	0.366	0.075	0.026	0.080	0.011	0.056	0.012	0.029	0.004	0.030	0.004	2.18	8.65	23.72	39.08	0.94	1.56	1.20
O1p-24-l-B	420	Black layer	0.939	1.839	0.224	0.784	0.161	0.046	0.128	0.020	0.118	0.025	0.076	0.011	0.071	0.012	4.45	8.65	24.62	37.50	0.93	1.51	0.97
O1p-24-b-B	420	Black band	0.724	1.250	0.138	0.520	0.109	0.024	0.094	0.016	0.096	0.016	0.045	0.007	0.048	0.006	3.09	8.44	25.51	44.84	0.91	1.10	1.10
O1p-37-b-B	355	Black band	0.977	1.783	0.217	0.830	0.152	0.026	0.141	0.021	0.116	0.023	0.066	0.012	0.062	0.009	4.44	8.84	26.05	42.10	0.89	0.83	1.15
O1p-58-n-B	315	Black nodule	0.508	0.811	0.123	0.479	0.097	0.017	0.081	0.012	0.077	0.015	0.043	0.009	0.068	0.010	2.35	6.44	33.19	34.93	0.75	0.92	0.54
O1–2y-b-B	140	Black band	0.648	1.100	0.137	0.498	0.078	0.073	0.086	0.014	0.077	0.034	0.047	0.008	0.050	0.007	2.86	7.84	12.72	19.22	0.85	4.17	0.96
O1–2y-n-B	140	Black nodule	0.721	1.045	0.161	0.673	0.128	0.042	0.135	0.021	0.128	0.021	0.065	0.012	0.096	0.012	3.26	5.66	42.07	33.55	0.71	1.51	0.55
O2d-l-B	18	Black layer	0.730	1.007	0.160	0.586	0.099	0.029	0.114	0.019	0.094	0.021	0.055	0.010	0.071	0.014	3.01	6.56	29.09	34.76	0.68	1.27	0.76
O2d-n-B	17	Black nodule	1.087	1.636	0.174	0.617	0.093	0.020	0.095	0.011	0.057	0.014	0.038	0.005	0.068	0.006	3.92	12.34	23.11	78.38	0.85	0.98	1.18
O1p-14-b-W	441	White band	0.312	0.577	0.069	0.248	0.052	0.011	0.043	0.008	0.041	0.008	0.026	0.004	0.028	0.006	1.43	7.74	38.75	39.42	0.91	1.06	0.83
O1p-24-n-W	420	White nodule	0.318	0.618	0.083	0.309	0.070	0.015	0.050	0.009	0.053	0.009	0.037	0.006	0.052	0.006	1.64	6.33	34.75	34.35	0.88	1.20	0.45
O1p-55-b-W	333	White band	0.304	0.473	0.069	0.227	0.053	0.014	0.052	0.009	0.050	0.012	0.031	0.007	0.044	0.005	1.35	5.45	32.52	26.26	0.75	1.28	0.51
O1p-58-n-W	323	White nodule	0.356	0.607	0.078	0.300	0.073	0.016	0.056	0.008	0.056	0.012	0.040	0.006	0.042	0.006	1.65	6.32	27.18	29.95	0.84	1.15	0.62
O1p-58-Q-W	315	Quartz grains	0.395	0.690	0.080	0.300	0.063	0.011	0.057	0.009	0.065	0.012	0.039	0.006	0.048	0.005	1.78	6.39	34.41	33.46	0.89	0.87	0.60

. The total REE concentration (∑REE) is generally low, ranging from 1.35 to 4.45 μg/g. PAAS-normalized REE patterns are weakly right-inclined for black siliceous rocks and weakly left-inclined for white varieties (Figure 4). The LREE/HREE ratios vary from 5.45 to 12.34. A negative Ce anomaly (δCe = 0.68–0.94) is observed, and most samples show a positive Eu anomaly (δEu = 0.83–4.17). The Y–Ho ratios range between 12.72 and 42.07 (mean 28.98), while La–Ho ratios for most samples range from 19.22 to 44.84 (mean 34.22).

4.2.3. Characteristics of Silicon and Oxygen Isotope

Silicon and oxygen isotope data are summarized in

Table 4. Si and O isotope data of the siliceous rocks from the Cement Plant Section.

Sample	Type	δ18O (V-SMOW)	δ30Si (V-NBS28)
		‰	‰
O1p-14-l-B	Black layer	26.19	2.84
O1p-14-b-B	Black band	25.79	3.28
O1p-24-b-B	Black layer	25.66	3.67
O1p-58-n-B	Black nodule	25.79	2.87
O2d-l-B	Black layer	29.33	2.79
O1p-24-n-W	White nodule	26.48	2.17
O1p-55-b-W	White band	26.44	2.42
O1p-58-n-W	White nodule	25.12	2.04
O1p-58-Q-W	Quartz grains	26.19	1.79

. The δ³⁰Si values are generally high, ranging from 2.79‰ to 3.67‰ (mean 3.09‰) in black siliceous rocks and from 1.79‰ to 2.42‰ (mean 2.11‰) in white varieties. Theδ¹⁸O_V-SMOW values range from 25.12‰ to 29.33‰, with a mean of 26.33‰.

4.3. Characteristics of Fluid Inclusion

Two types of fluid inclusions are identified within the siliceous rocks. The first type is liquid-rich inclusions, which are small in number and occur in clusters within the microcrystalline quartz of black siliceous rocks from the Penglaiba Fm and Dawangou Fm. The second type is hydrocarbon-bearing aqueous inclusions (gas–liquid, two-phase), which are distributed in banded and zonal patterns within and along the margins of third-generation phanerocrystalline quartz in white siliceous rocks of the Penglaiba Fm. As shown in

Table 5. Characteristics of homogenization temperatures and salinities of fluid inclusions in quartz from the Cement Plant Section.

Sample	Type		Size	Th	Salinity
			μm	°C	wt% Nacl
O1p-14-l-B	Aqueous-rich inclusion	Quantity	6	6	6
		Min	2 × 4	150	16.55
		Max	6 × 12	162	16.63
		Mean		155.5	16.58
O2d-l-B	Aqueous-rich inclusion	Quantity	7	7	7
		Min	1 × 5	139	14.21
		Max	4 × 4	150	14.36
		Mean		144.3	14.31
O1p-58-Q-W	Hydrocarbon-aqueous inclusion	Quantity	23	23	23
		Min	1 × 3	100	18.63
		Max	5 × 7	126	20.15
		Mean		111.1	19.4

, the liquid-rich inclusions occur in clusters within the chert (Figure 5a). In the Dawangou Fm, they measure 1–4 μm in length and 4–5 μm in width, with homogenization temperatures of 139–150 °C (Figure 5d) and salinities of 14.21%–14.36% (Figure 5e). In the Penglaiba Fm, they are 2–6 μm in length and 6–12 μm in width, with homogenization temperatures of 150–162 °C and salinities of 16.55%–16.63%. The hydrocarbon-bearing aqueous inclusions are distributed in bands within and along the edges of the third-generation quartz in white siliceous rocks (Figure 5b,c). They measure 1–3 μm in length and 5–7 μm in width, with homogenization temperatures of 100–126 °C (Figure 5d) and salinities of 18.63%–20.15% (Figure 5e).

4.4. Reservoir Characteristics of Siliceous Rocks and Their Host Rocks

In the study area, the main reservoir spaces in siliceous rocks (Penglaiba to Dawangou Fm) are intercrystalline pores, intraparticle pores, and fractures, while in hydrothermally influenced carbonate host rocks they are dissolution pores, interparticle pores, and fractures.

Pores in siliceous rocks are barely visible under optical microscopy (Figure 3a–d). Under an electron microscope, intercrystalline pores, coexisting with microcrystalline quartz, are micron-sized with pore diameters of approximately 2.5–5 μm (Figure 6a,b). Compared to dissolution pores in carbonate rocks, the intercrystalline pores are relatively uniform. Intraparticle dissolution pores, mainly in the granular quartz of white siliceous rocks, are about 0.1–0.3 μm in size (Figure 6c) and are less developed than intercrystalline pores.

Fractures constitute important reservoir spaces in both siliceous rocks and host rocks, exhibiting a variety of types, including straight bedding-parallel fractures, low-angle pressure solution seams, high-angle shear fractures, and crystal-boundary fractures between quartz mineral grains. In black and white siliceous rocks, fractures often form intersecting networks and vary from unfilled to fully filled, mainly with calcite or quartz (Figure 6d–f). In black siliceous rocks, fractures 2–5 mm wide cut cryptocrystalline-microcrystalline quartz and are filled with calcite or quartz (Figure 6f). White siliceous rocks contain quartz grain-boundary fractures that are 0.01–0.03 mm wide (Figure 6e). In hydrothermally altered medium-crystalline dolostone, saddle dolomite shows dark red cathodoluminescence (Figure 6f,g). The reservoir space in this dolostone is dominated by intercrystalline pores, with pore sizes of about 10–30 μm. In crystalline limestone, fractures are the main reservoir space (Figure 6h).

Porosity and permeability measurements and statistics for different types of siliceous rocks from Penglaiba, Yingshan, and Dawangou Fm are summarized in

Table 6. Porosity and permeability data of Middle–Lower Ordovician siliceous rocks and their host rocks from the Cement Plant Section, Kalpin area.

Reservoir Properties	Black Siliceous Rock	White Siliceous Rock	Medium-Crystalline Dolomite	Fine-Crystalline Dolomite	Crystalline Limestone	Calcarenite	Bioclastic Micritic Limestone	Micritic Limestone
Maximum porosity (%)	4.88	12.82	7.87	3.33	4.63	4.61	3.93	0.80
Minimum porosity (%)	1.68	5.58	2.51	1.02	1.97	1.48	0.76	0.36
Mean porosity (%)	3.41	8.82	5.17	2.28	3.62	2.30	2.29	0.56
Maximum Permeability (mD)	3.237	25.18	1.902	0.034	0.935	0.741	0.197	0.002
Minimum Permeability (mD)	0.009	0.350	0.002	0.001	0.010	0.001	0.001	0.001
Mean Permeability (mD)	0.747	6.911	0.495	0.005	0.222	0.087	0.032	0.001

. Black siliceous rocks have a porosity of 1.68%–4.88% (mean 3.41%) and permeability of 0.009–3.237 mD (mean 0.747 mD). White siliceous rocks show higher porosity (5.58%–12.82%; mean 8.82%) and permeability (0.35–25.18 mD; mean 6.911 mD). Hydrothermally altered medium-crystalline dolostones adjacent to siliceous rocks have a porosity of 2.51%–7.87% (mean 5.17%) and permeability of 0.002–1.902 mD (mean 0.495 mD). Altered crystalline limestones exhibit a porosity of 1.97%–4.63% (mean 3.62%) and permeability of 0.010–0.935 mD (mean 0.222 mD). In contrast, rocks unaffected by hydrothermal activity have a lower porosity (0.36%–4.60%; mean 2.03%) and permeability (0.001–0.74 mD; mean 0.04 mD).

5. Discussion

5.1. Source of Siliceous Fluids

Geochemical data indicate two distinct diagenetic fluids for the Penglaiba–Dawangou siliceous rocks. Black siliceous rocks were formed by a mixed hydrothermal fluid composed of deep-sourced magma that had dissolved Precambrian basement clastic rocks, formation brines, and seawater. White siliceous rocks originated from a mixed hydrothermal fluid of deep Si-rich fluids (heated at depth) and formation brines.

5.1.1. Major and Trace Elements

Major and trace element compositions confirm a hydrothermal origin for all studied siliceous rocks. Black varieties formed from a mixture of deep (lower crust/mantle) hydrothermal fluid and seawater, whereas white varieties likely precipitated from deeper formation water hydrothermal fluids. Enrichment of Al and Ti reflects the input of terrigenous aluminosilicates, whereas enrichment of Fe and Mn is primarily associated with hydrothermal activity [40]. Consequently, the Al–Fe–Mn diagram is commonly used to assess the contributions of terrigenous clastics and hydrothermal inputs to sediments [41,42,43]. The low Al and Ti but high Fe and Mn in our samples point to weak terrigenous influence and a strong hydrothermal association. All samples were plotted within the hydrothermal field on an Al–Fe–Mn diagram (Figure 7a), confirming their hydrothermal origin.

The correlation between Cr, Ni, and Fe can indicate the involvement of mantle-derived hydrothermal fluids [44]. Black siliceous rocks show a positive Cr–Fe and Ni–Fe correlation, unlike white varieties (Figure 7b,c), suggesting that mantle-derived fluids contributed Fe to the black rocks but not dominantly to the white rocks. Y–Ho and Th–U ratios help constrain fluid sources and diagenetic settings [45,46,47,48]. In the Cement Plant Section, black siliceous rocks have Th–U values less than 1 (slightly higher in Yingshan and Dawangou due to seawater influence) and Y–Ho ratios close to the chondritic value (≈28), plotting in the metasomatic field (Figure 8a). This indicates a fluid mixture of deep (lower crust/mantle) hydrothermal fluid and seawater [19,49]. Black siliceous rocks from well KP1 have both Th–U (0.015–0.018) and Y–Ho (15–18) ratios lower than those from the Cement Plant Section samples (Figure 8a), representing typical characteristics of deep hydrothermal fluids [19,47,49] and possibly indicating less seawater influence in this well area. White siliceous rocks from the Cement Plant Section have Th–U values between deep hydrothermal and seawater end members and high Y–Ho (32.52–38.75). Their data are plotted in metasomatic and formation water–hydrothermal mixing fields; white rocks from KP1 are also plotted in the metasomatic field (Figure 8a), suggesting a diagenetic fluid of relatively deep formation water–hydrothermal origin. The covariation between La–Ho and Y–Ho indicates that both rock types experienced a high degree of silicification (Figure 8b), reflecting prolonged hydrothermal migration and extensive water–rock reactions [47].

5.1.2. Rare Earth Elements

The rare earth element (REE) characteristics support a hydrothermal origin for the siliceous rocks. Both black and white varieties have low ΣREE (<200 × 10⁻⁶), which is typical of hydrothermal systems [50]. In the Cement Plant Section and KP1 well, layered and banded black siliceous rocks from the Penglaiba Fm and banded black siliceous rocks from the Yingshan Fm show right-inclined to flat REE patterns (La_N/Yb_N = 0.96–1.20; Figure 4a) and pronounced positive δEu anomalies (0.98–4.17; Table 3), which are hallmark hydrothermal features [51]. In contrast, other types in the section (e.g., black nodules from Penglaiba, black layers from Dawangou, and all white siliceous rocks) display weakly left-inclined patterns (La_N/Yb_N = 0.45–0.91; Figure 4a,b) with subduedδEu anomalies (0.87–1.28), indicating a weaker hydrothermal overprint. Among them, the layered siliceous rocks have δEu values 0.2–0.5 higher than other types, indicating a stronger hydrothermal signature. All types of siliceous rocks in the section show weak negative δCe anomalies (0.68–0.94), differing from mantle or crustal signatures. In contrast, both black and white rocks from the KP1 well show near-normal δCe (1.02–1.07), suggesting that hydrothermal fluids at the section were diluted by seawater or formation water, whereas fluids in the KP1 well area experienced less dilution [52,53]. The δEu values of the main samples are lower than those of high-temperature hydrothermal minerals like sinter (average δEu: 2.15) [53], further indicating a genesis involving the mixing of low-temperature hydrothermal fluids and seawater. This seawater influence is particularly evident in the black siliceous rocks of the Dawangou Fm, which show more pronounced negative δCe and weaker positive δEu anomalies (Table 3, Figure 4a).

5.1.3. Silicon and Oxygen Isotopes

The δ³⁰Si of siliceous rocks is unaffected by diagenesis and late-stage fluids, making it an effective tracer for their diagenetic environment [54]. Compared to previous studies on siliceous rocks of different origins [2,23,49,54,55,56,57,58,59,60,61,62,63] (Figure 9), the δ³⁰Si values of the black chert samples in the study area range from 2.79‰ to 3.67‰, and those of the white siliceous rocks range from 1.79‰ to 2.42‰. These values are similar to those reported for siliceous rocks in the Penglaiba Fm from the KP1 well in this region [20] (Figure 9). Furthermore, in previous studies concerning the nature of siliceous hydrothermal fluids in the Tarim Basin, the average δ³⁰Si value in the Gucheng area was 2.5‰ [15], while δ³⁰Si values ranged from 2.1‰ to 3.0‰ in chert from the Shunnan area and from 2.3‰ to 2.7‰ in quartz from the veins [17]. This consistency suggests that the siliceous fluids in the Tarim Basin likely share a similar origin and were characterized by episodic activity. Critically, they are significantly higher than typical deep hydrothermal or seawater values, aligning with a metasomatic diagenetic origin (Figure 9a). The δ¹⁸O_V-SMOW values of both types of siliceous rocks (25.12‰–29.33‰) also align with the characteristics of diagenetic quartz (Figure 9b). Therefore, δ³⁰Si and δ¹⁸O indicate a diagenetic–metasomatic origin for the siliceous rocks.

This interpretation is consistent with petrographic observations: the absence of biogenic skeletons and volcanic hydrothermal minerals (e.g., fluorite, barite), and the presence of clear hydrothermal replacement and precipitation textures. Consequently, the diagenetic fluid for the siliceous rocks was likely a mixture of deep metamorphic fluid and seawater (Figure 9a). Elevated δ³⁰Si values can be driven by a high water–rock ratio and frequent hydrothermal activity [2,64]. Therefore, the higher δ³⁰Si in black siliceous rocks suggests formation in a high water–rock ratio, open, and shallow-burial environment with strong seawater influence. In contrast, the lower values in white varieties point to formation in a low water–rock ratio, closed system during medium to deep burial.

5.1.4. Fluid Inclusions

Fluid inclusion data indicate that the diagenetic temperature of black siliceous rocks (139–162 °C) is higher than that of the white siliceous rocks (100–126 °C), with both falling within the low-temperature hydrothermal range. The diagenetic temperature of the black siliceous rocks exceeds the maximum burial temperature (100–120 °C) of the Middle–Lower Ordovician in the study area [65,66,67]. In contrast, the diagenetic temperature of the third-generation quartz, formed during the final stage of the diagenetic fluid in the white siliceous rocks, is comparable to the maximum burial temperature, supporting an allochthonous (externally derived) origin for the silica-bearing fluids. The salinities of fluid inclusions are high: 14.23%–16.63% for those in microcrystalline quartz (Q1) and 18.63%–20.15% for those in phanerocrystalline quartz (Q3). These values significantly exceed Ordovician seawater salinity (5.5%) and the salinity of local magmatic–hydrothermal fluorite (7%) but match the range of distal deep hydrothermal fluids (15%–25%) driven by magmatic activity [68,69], indicating that the siliceous fluids were not derived from a single seawater or magmatic source. The salinity of fluid inclusions in the sub-salt Cambrian dolomite in the study area (6%–10%) is also significantly lower than that in the Ordovician siliceous rocks [68]. Combined with elemental geochemistry, these results indicate that the moderate salinity in black siliceous rocks reflects dilution by seawater, while the high salinity in white siliceous rocks points to a dominant contribution from deep formation brines with limited seawater influence. Comparative analysis with previous studies reveals that in the Tanggu Depression of the Tazhong area, fluid inclusions within quartz veins hosted by the Yingshan Fm limestone exhibit homogenization temperatures ranging from 125 to 155 °C, with salinities around 17.5% [21]. Similarly, in the Shunnan area, cherts from the Yingshan to Yijianfang Fm show homogenization temperatures of 143–170 °C and salinities of approximately 21.5% [19]. These indicate that the diagenetic fluids responsible for siliceous rock formation in the Kalpin area share similarities in temperature and salinity with those in other parts of the Tarim Basin, suggesting a potentially common source. However, variations in hydrothermal temperature and the resulting impact on carbonate reservoirs are still evident, likely controlled by local structural differences and proximity to heat sources.

5.2. Timing of Siliceous Rock Formation

The absence of precise dating methods for siliceous rocks makes it difficult to determine their formation age. Consequently, timing must be inferred from an integrated analysis of fluid inclusions, stratigraphic relationships, and regional tectonic burial–thermal history. As established, the siliceous fluids are deep-sourced, suggesting that tectonic movements and heat flow are crucial for both mineralization stages. Situated at the southern Tianshan front, the Kalpin area has experienced intense tectonic activity since the Precambrian [28,29,30]. Major faulting episodes occurred at 450–420, 370–300, 200–100, 85–65, and 40–0 Ma [67], with hydrothermal pulses mainly in the Middle–Late Caledonian [70,71,72], Early Hercynian [70,71,72,73], and Late Hercynian [39,74]. Consequently, the Middle Ordovician to Permian interval provided the necessary fault and hydrothermal conditions for silicification.

The hydrocarbon-bearing aqueous inclusions in third-generation quartz of white siliceous rocks are key chronological indicators. Integrating the reservoir formation history reconstruction of the Ketan-1 well in the area and Re–Os isotopic dating of hydrocarbons, the main hydrocarbon charging periods for the Cambrian source rocks in the study area were the Middle–Late Devonian (386–366 Ma) and the Eocene (24–30 Ma) [75]. Projecting the temperature of these inclusions (average 111 °C) onto the burial history curve corresponds to the Late Devonian and Permian–Triassic periods (Figure 10). Because Ordovician were already exposed during the low-geothermal gradient Eocene, and because these inclusions reside in the lowest-temperature quartz at fracture cores, they must have precipitated late in the fluid evolution. Therefore, the white siliceous rocks mainly formed in the Middle–Late Devonian (~386–362 Ma). This period coincides with the intra-oceanic subduction of the South Tianshan Ocean, deep hydrothermal upwelling from ~370 Ma [73], and the onset of Early Hercynian faulting [67,71], which together supplied the requisite heat and fluid pathways. Hence, the white siliceous rocks are dated to 370–362 Ma.

For black siliceous rocks, which lack hydrocarbon-bearing inclusions, timing relies on field and petrographic relations. Although some earlier studies suggested a Permian [15,18,22], field evidence shows that black siliceous rocks are cut and filled by white varieties (Figure 2d–f) and later shear fractures, proving they formed earlier. Petrography reveals black chert replacing calcarenite and dolostone, with relict dolomite (Figure 2a), indicating that silicification post-dated early dolomitization. Previous studies suggest that shallow-burial dolomitization in the Kalpin area primarily occurred from the Late–Middle Ordovician to the Early Silurian [76], which constrains the minimum age (not earlier than the end of the Middle Ordovician) for the formation of black siliceous rocks. Their macroscopic field characteristics, including extensive lateral distribution along bedding planes, replacement of early dolostone, and geochemical signatures indicative of seawater influence, all suggest formation during the early diagenetic stage (Middle–Late Ordovician). Regional tectonic studies also support this inference: Late Ordovician subduction of the South Tianshan Ocean triggered intense magmatism [70,71,72,77] and the development of deep faults (e.g., ancient Tumuxiuke, Shajingzi) that controlled sedimentation in the NW basin (Wushi–Kalpin–Bachu) [78,79]. Recent seismic data (Kenan-1 well) and outcrop studies (Middle–Late Caledonian fractures at the Cement Plant Section) confirm a fault system penetrating Middle–Lower Ordovician near the Ordovician–Silurian boundary [27,80,81], providing conduits for deep fluid upwelling. Moreover, the absence of siliceous rocks in the Upper Ordovician Qilang and Yingan limestones precludes a Permian fault-controlled origin. In summary, the black siliceous rocks formed during early diagenesis from the Late–Middle Ordovician to pre-Late Devonian, closely linked to the Late Ordovician (450–440 Ma) tectonic–thermal event. The white siliceous rocks formed later in the Middle–Late Devonian (370–362 Ma).

5.3. Diagenetic Model of Siliceous Rocks

During the Cambrian–Ordovician, the Tarim Basin was a stable cratonic depression, developing marine basin, continental shelf, and platform. By the Late Ordovician, influenced by the subduction of the Paleo–Tianshan Ocean, intense tectonic activity occurred along the northern margin of the basin, forming numerous deep, large-scale strike-slip fault zones that penetrated the basement and the Ordovician [78,79]. Accompanying plate movements, magmatic hydrothermal fluids rich in elements such as Si, Fe, and Cr, originating from the lower crust or mantle, migrated upward along these deep faults. While flowing through Precambrian basement clastic rocks, the high-temperature hydrothermal fluids dissolved siliceous clasts around the fractures. Subsequently, passing through the Cambrian evaporite layers, they mixed with high-salinity formation fluids within these successions, resulting in a significant increase in the salinity of the mixed hydrothermal fluid. When the fluids entered the Ordovician carbonate formations, Penglaiba to Dawangou Fm were in a shallow burial environment of approximately 200–600 m. As the faults had already penetrated the Ordovician, seawater could infiltrate downward through the fractures and mix with the upwelling hydrothermal fluids. This process lowered the salinity and temperature of the hydrothermal fluid, ultimately forming a medium-salinity (14.23%–16.63%) and relatively high-temperature (139–162 °C) hydrothermal fluid derived from a mixture of magmatic, formation brine, and seawater sources. During this period, compaction in the Penglaiba–Dawangou Formations was weak, and pores and fractures were relatively well-developed, allowing the hydrothermal fluids to migrate laterally along carbonate bedding planes and fault zones (Figure 11a) [82]. Through intense water–rock interaction, the siliceous hydrothermal fluids dissolved and replaced the carbonate rocks, preserving the granular ghost textures of the original carbonates within the siliceous rocks (Figure 3b). The CO₂ generated during dissolution dissolved into the fluid, causing the hydrothermal solution to quickly become oversaturated with SiO₂ and weakly acidic, further dissolving carbonates to create enlarged pores and fractures. This also promoted the rapid precipitation of silica, forming abundant chert. In the Penglaiba Fm, fluctuations in sea level led to variations in sedimentation rates, resulting in the interbedded deposition of thin and thick dolomite layers. Due to mechanical differences, fractures were more likely to form between thin layers, and the crystalline texture of the dolomite provided permeable pathways for hydrothermal fluid intrusion. Consequently, extensively distributed thin-bedded or banded siliceous rocks formed between thin dolomite layers and at interfaces between thin and thick layers near the faults (Figure 11b). As lateral migration distance increased and temperatures gradually decreased, the siliceous rocks transitioned to predominantly discontinuous bands and nodules (Figure 11b). When the hydrothermal fluids migrated upward into the Yingshan and Dawangou Fm, their vertical and lateral migration within the limestone was somewhat restricted by the relatively dense micritic limestone developed in the middle sections. However, longer-distance lateral migration occurred within the well-porous and fractured sparite grainstone and at lithological interfaces between grainstone and micritic limestone, forming bedded siliceous rocks. Within the micritic limestone, siliceous rocks mostly occurred as nodules along fracture zones (Figure 11c). Overall, this episode of silica-rich fluid was relatively high in Fe content, resulting in predominantly black-colored siliceous rocks. As the carbonates were in a high water–rock ratio, open diagenetic environment during shallow burial, water–rock reactions were thorough and prolonged. The black siliceous rocks exhibit high Th–U ratios and relatively high δ³⁰Si values (avg. 3.99).

By the Middle to Late Devonian, the Ordovician had entered a medium to deep burial stage (2500–3500 m) [66,67]. Influenced by the subduction of the South Tianshan Ocean, tectonic activity in the study area revived, triggering a new phase of hydrothermal event [72,73]. The heat source was primarily located in the southwestern depression, Bachu uplift, and northern Tarim uplift areas [74]. Therefore, the hydrothermal fluids in the Kalpin area were mainly derived from the distal heating of Precambrian basement formation fluids by magmatic activity, which drove the upward migration of silica-rich fluids. Compared to the first episode, the mineralization temperature of this silica-rich fluid was relatively lower (avg. 111.1 °C), and due to the lack of elements such as Fe and Cr, which are relatively enriched in the mantle, the resulting siliceous rocks appeared grayish-white. The hydrothermal fluids mixed with Cambrian brines as they migrated upward and then entered the Middle–Lower Ordovician carbonates (Figure 11d). Compared to the shallow burial stage, the lack of seawater dilution during deep burial and the formation of quartz minerals in a relatively closed environment resulted in lower δ³⁰Si values (avg. 2.11) and higher salinity of fluid inclusions trapped during mineral formation (avg. 19.4%) in the siliceous rocks formed by this fluid episode. By this time, the Ordovician had become highly compacted due to pressure dissolution and cementation. Effective reservoir space in the Penglaiba Fm mainly consisted of intercrystalline pores in medium-crystalline dolomite, fractures at interfaces between early-stage siliceous rocks and carbonates, and newly formed fractures. The dense micritic limestone at the top of the Penglaiba Fm and within the Yingshan Fm restricted the vertical flow of hydrothermal fluids (Figure 11f). Consequently, the siliceous fluids slowly seeped through the confined fracture networks in the middle and lower dolomite sections of the Penglaiba Fm (Figure 11e). Under conditions of low temperature and low flow rate, the fluids slowly dissolved and replaced carbonate minerals along fractures, generating minor amounts of replacement-derived microcrystalline quartz. With further cooling, chalcedony → columnar → granular quartz sequentially precipitated in residual pores and fractures (Figure 3e–g), ultimately forming white-banded and nodular siliceous rocks. Within the Yingshan–Dawangou limestones, due to the blockage by the aforementioned dense barriers and the increased compactness of limestone during medium-deep burial, hydrothermal fluids struggled to migrate upward or laterally. Therefore, white siliceous rocks are scarcely developed within these limestones (Figure 11f).

5.4. Significance of Siliceous Rocks for Reservoir Formation

The effectiveness of siliceous rock reservoirs is jointly controlled by their petrophysical properties (porosity and permeability) and development scale (spatial distribution and continuity) [83]. These two aspects are fundamentally constrained by differential diagenetic evolution and multi-stage fluid–rock interactions. The results of this study indicate that the impact of siliceous fluids on the reservoir properties of carbonate rocks is not a simplistic process of “early-stage porosity enhancement and late-stage pore blockage”; the actual relationship is more complex. Based on the analysis of reservoir space and petrophysical characteristics of both siliceous rocks and carbonate reservoirs, combined with the diagenetic model of siliceous rocks, this paper discusses the influence of the mechanisms of two episodes of siliceous diagenetic fluids on reservoir formation.

The black siliceous rocks formed during the late Caledonian episode developed during the shallow burial stage of the Penglaiba Fm to Dawangou Fm. At this stage, carbonates experienced weak compaction, and primary pores and fractures were relatively well-developed, especially within grainstone and dolomite intervals, providing favorable pathways for hydrothermal fluid invasion. The first stage of high-temperature, silica-rich hydrothermal fluids entered the carbonates along bedding planes, lithological interfaces, and fracture zones, dissolving minerals such as calcite and dolomite. The inorganic CO₂ generated during dissolution re-dissolved into the fluid as the temperature decreased, significantly increasing fluid acidity and thereby promoting continuous dissolution [17]. This process created a large number of enlarged dissolution fractures and vugs. Simultaneously, SiO₂ rapidly reached saturation and precipitated, replacing carbonate minerals in the form of cryptocrystalline to microcrystalline quartz or chalcedony (Q1) (Figure 3a–c), partially filling the dissolved spaces (Figure 3d–i) and forming layered, bedding-parallel bands and nodular black siliceous rocks (Figure 2a–e). These siliceous rocks have a dense structure, with reservoir space dominated by micron-scale intergranular pores, resulting in low porosity (avg. 3.41%; Table 6; Figure 12), which also led to a decrease in the porosity of the carbonate reservoir. However, silicification also inhibited the destruction of reservoir space by later compaction and cementation. Furthermore, the high brittleness of the siliceous rocks, combined with their early extensive lateral distribution and vertically stacked multi-layer nature [84], formed a certain scale of brittle siliceous rock bodies, laying the groundwork for subsequent structural modification. Subsequent multi-phase tectonic movements (Hercynian, Indosinian, and Himalayan) caused the development of dense networks of shear–tensile fractures within the early formed siliceous rocks (Figure 6d–f). These fractures significantly improved permeability. The permeability of black siliceous rocks (avg. 0.747 mD) can be two orders of magnitude higher than that of unsilicified dense micritic limestone (avg. 0.001 mD) and fine-crystalline dolomite (avg. 0.005 mD) (Table 6, Figure 12A–D), making them preferred pathways for hydrocarbon migration. The micro-intergranular pores developed within them also contribute to some storage space (Figure 6a–c). Exploration examples, both domestically and internationally, confirm that such siliceous rocks can form effective reservoirs (Figure 13). For instance, in the Parkland oil field in Canada, approximately 18 m-thick siliceous rocks near a strike-slip fault zone within the Wabamun Group carbonates are densely fractured, exhibiting porosity as high as 27% and permeability ranging from 1 to 800 mD, supporting high-production development [11]. In the Tarim Basin, a 10 m-thick siliceous rock interval in the Penglaiba Fm of the TS-6 well in the Tahe oil field has an average porosity of 14.74% and permeability of 0.331 mD, constituting an effective gas-producing reservoir [16]. An analysis of the petrophysical properties of siliceous rocks in the Yingshan Fm of the Shunnan area [85] and the Middle–Lower Ordovician of the Kalpin area shows that even when the porosity of siliceous rocks is low (<5%), the development of fractures within them can still significantly enhance their permeability (Figure 12).

In addition to the inherent reservoir potential of the siliceous rocks themselves, the high temperature of this hydrothermal episode induced thermal alteration and recrystallization of the surrounding carbonates. This resulted in the formation of medium- to coarse-crystalline dolomite (Figure 12A,B) and crystalline limestone (Figure 12C,D) adjacent to the layered black siliceous rocks, accompanied by the development of intercrystalline pores (10–50 μm in diameter) and intercrystalline fissures (Figure 6g,i). In contrast, the carbonate rocks surrounding banded or nodular siliceous rocks, having undergone weaker hydrothermal alteration, largely retain their original fine-crystalline dolomite or sparry grainstone fabric (Figure 12A–D). This suggests that larger-scale siliceous rocks are more likely to exert significant thermal alteration effects on their host rocks. Compared to unaltered fine-crystalline dolomite, the medium-crystalline dolomite exhibits a 3%–7% increase in porosity and a permeability enhancement of 100–1000 times (Table 6; Figure 12A,B). Similarly, the crystalline limestone shows a 2.5%–3.8% increase in porosity and a permeability improvement of 10–200 times over the original limestone (Table 6; Figure 12C,D). This confirms a significant enhancement in the petrophysical properties of the altered wall rocks, demonstrating a positive impact of siliceous hydrothermal activity on carbonate reservoirs. Furthermore, the scale of hydrothermal carbonate development appears to increase with the scale of the associated siliceous rocks. Comparisons with other exploration areas within the basin and global analogs reveal a consistent pattern (Figure 13): a 20 m-thick siliceous rock interval in the Parkland oil field is associated with a 20 m-thick hydrothermal dolomite reservoir [11]; an 18 m-thick siliceous rock in the Tahe oil field is surrounded by a 15 m-thick alteration zone, whereas a 3.75 m-thick siliceous body is flanked by only a 1.5 m-thick dolomite reservoir [16]; in the SN4 well, a 1 m-thick siliceous layer in the Yingshan Fm has altered its host rock over a distance of less than 0.3 m [85] (Figure 13); similarly, thin-bedded siliceous rocks (5–20 cm) in the Cement Plant Section have only affected the surrounding rock within a narrow 10–20 cm halo (Figure 12A,B). This relationship indicates that dolomite, due to its inherent propensity to develop fractures and intercrystalline pores, is more susceptible to hydrothermal alteration than limestone. Moreover, thicker, laterally continuous layered siliceous rocks within dolomitic sequences serve as more reliable indicators for exploring hydrothermal reservoirs. Silica-rich hydrothermal fluids can promote recrystallization and pore–fracture development in the host rock, creating a composite siliceous–carbonate reservoir system that substantially expands the scale of effective hydrothermal reservoirs.

In contrast, the second episode of silica-rich fluid, which formed the white siliceous rocks, occurred during a medium-to-deep burial stage (approximately 3000 m in depth). By this time, the Ordovician carbonates had become extremely compact due to intense compaction and extensive cementation, leading to a near cessation of material exchange. Although the Hercynian orogeny generated numerous deep faults, allowing the relatively low-temperature silica-rich fluids of this second episode to infiltrate the carbonate reservoirs along fractures (Figure 12D), the lack of effective flow pathways within the highly compacted rock matrix resulted in slow fluid migration. Consequently, as the fluid cooled gradually during migration, silica precipitated as micro- to coarse-crystalline quartz over multiple generations, forming only thin-banded (often <1 m) or nodular white siliceous rocks within major fracture systems. The quartz grains in these rocks are coarser than the cryptocrystalline quartz and chalcedony precipitated rapidly during the Late Caledonian episode. This coarser texture results in more developed intercrystalline pores and intergranular fractures, giving these rocks high porosity and high permeability (averages of 8.82% and 6.91 mD, respectively; Table 6; Figure 12A,B). However, the lower fluid temperature during their formation resulted in a weaker thermal alteration effect and minimal modification of the surrounding reservoir rock, failing to induce recrystallization in the adjacent dolomite (Figure 12B). Furthermore, despite their excellent intrinsic porosity and permeability, the white siliceous rocks are limited in scale and lateral continuity, making it difficult for them to form large-scale effective reservoirs. Overall, their contribution to improving reservoir quality is negligible.

In summary, the first stage of silica-rich fluid had a more significant impact on modifying the carbonate reservoir compared to the second episode. The early fluid invaded through faults, early diagenetic interfaces, bedding planes, and pore–fracture networks within high-energy facies. It both dissolved and enlarged the original reservoir space and subsequently filled it rapidly, reducing primary porosity but creating brittle, silica-rich zones. These zones later became preferred sites for stress release during subsequent tectonic events. Multiple tectonic activities throughout geological history reactivated or generated new fractures within these zones, forming dense fracture networks and micro-pores that reshaped the reservoir’s flow pathways and storage capacity. Therefore, an effective reservoir formation model can be summarized as “early silicification establishes the brittle foundation, while later tectonic fracturing achieves reservoir modification.” Future exploration should focus on favorable facies zones where large-scale siliceous rocks (providing the brittle material basis) coincide with areas of well-developed late-stage structural fractures.

6. Conclusions

Based on detailed petrographic, microthermometric, and geochemical investigations of Middle–Lower Ordovician siliceous rocks from the Cement Plant Section in the Kalpin area, northwestern Tarim Basin, the following conclusions can be drawn:

• The Middle–Lower Ordovician in the Kalpin area of the northwestern Tarim Basin hosts two stages of siliceous rocks: black and white. The black siliceous rocks are of hydrothermal replacement origin. The white siliceous rocks result from both hydrothermal replacement and hydrothermal precipitation.
• The black siliceous rocks formed during the Late Ordovician, sourced from siliceous fluids comprising a mixture of deep mantle-derived components, basaltic clastics, high-salinity formation brines, and shallow seawater, characterized by relatively high temperatures and SiO₂ supersaturation. The white siliceous rocks formed during the Late Devonian (Early Hercynian period), sourced from silicon-rich fluids mixing basaltic clastics and high-salinity formation brines, characterized by lower temperatures and slightly lower silica content.
• The black siliceous rocks developed on a larger scale with multiple layers and are susceptible to later tectonic modification, which generates fracture networks. This can increase their permeability by more than two orders of magnitude, forming fracture porosity reservoirs with considerable potential.
• The black siliceous rocks induce more pronounced hydrothermal alteration in the surrounding rocks. The vertical alteration zone of a single layer ranges from 10 to 20 cm in thickness and expands with increasing layer thickness or cumulative thickness. In dolomite sequences, continuous, laterally stable, and thick (10–20 m) black siliceous rock bodies, together with the adjacent medium-crystalline dolostones altered by hydrothermal fluids, represent more promising hydrothermal reservoir targets in the study area.