Origin and Reservoir Significance of Authigenic Minerals in Lacustrine Shales: A Case Study from the Paleogene Dongying Sag, Bohai Bay Basin, East China

Abstract

Authigenic minerals in shale are products of the co-evolution of organic and inorganic components, affecting the heterogeneity of shale reservoirs. However, due to their fine granularity and complex rock composition, studies on these minerals in shale are still insufficient. This research focuses on the lacustrine shales from the upper sub-member of the fourth member in the Eocene Shahejie Formation, Dongying Sag, East China. Utilizing core samples, thin sections, scanning electron microscope, X-ray diffraction, elemental geochemistry, and organic geochemistry, we systematically characterized the features and origins of authigenic minerals. The results identified several typical authigenic minerals, including authigenic quartz, framboidal and euhedral pyrite, ferroan dolomite, kaolinite, chlorite, and albite. Authigenic quartz is predominantly diagenetic silica formed through smectite illitization, acidic dissolution of K-feldspar, and alkaline dissolution of detrital quartz. Pyrite is a product of microbial sulfate reduction, with framboidal pyrite forming during an early diagenetic stage under conditions with sufficient solute supply and euhedral pyrite forming during a later stage under conditions with insufficient solute supply. Ferroan dolomite originates from the precipitation of Fe and Mg during smectite illitization, with slight contributions from the acidic dissolution of chlorite and calcite. Kaolinite stems from the acidic dissolution of K-feldspar, while chlorite results from the transformation of kaolinite. Albite primarily arises from the alkaline alteration of anorthite and K-feldspar. Most non-clay authigenic minerals likely enhance reservoir quality by slightly reducing the effects of compaction, whereas authigenic clay minerals typically exert detrimental effects on reservoir properties. This study constrains the genesis of authigenic minerals to assess their influence on reservoir quality in lacustrine shale.

1. Introduction

Research on the diagenesis of clastic rocks has evolved over the past few decades from early qualitative observations to quantitative modeling, detailed predictions, and assessments of reservoirs [1,2,3,4,5,6]. This shift has significantly advanced the exploration of oil basins into more sophisticated domains [7,8]. Compared to clastic rocks and carbonate, organic-rich shales display similar detrital and authigenic minerals, but they are characterized by finer grain sizes, diverse compositions, and coexistence with organic matter [9,10,11,12]. Shale diagenesis is governed by both external controls (e.g., tectonic activity, depositional environment, burial history) and internal factors (e.g., lithofacies heterogeneity, organic matter composition and abundance, thermal maturity) [13,14,15,16,17]. The interplay between organic and inorganic diagenetic processes drives the formation of authigenic mineral assemblages, including microquartz, pyrite, sparry calcite, albite, clay minerals, and zeolites, which fundamentally shape reservoir properties [18,19,20]. Systematic analysis of these minerals provides key insights into microscale diagenetic systems and their control on the development of fine-grained reservoirs [6,14,21,22,23,24,25].

Extensive research has focused on diagenesis in organic-rich shales and its contributions to reservoir quality; however, relatively less attention has been paid to authigenic minerals formed during burial, despite their non-negligible impact on reservoir properties [26,27,28,29,30]. The formation of authigenic minerals is fundamentally controlled by diagenetic fluids. In the relatively closed diagenetic system of shale, the diagenetic fluids are primarily composed of connate sedimentary water, hydrocarbons, and acidic fluids released during the thermal maturation of organic matter, as well as interlayer water released during mineral transformation [9,11,31,32]. Early diagenetic processes in organic-rich shales are largely influenced by inherited primary sedimentary water, which has abundant bicarbonate ions, reduced manganese and iron, and sulfide [33,34,35]. They can be cemented into early diagenetic minerals such as calcite, pyrite, siderite, and dolomite [36,37,38,39]. During shallow burial, hydrocarbons are metabolized by biological processes and redox reactions, such as those involving methanogenic bacteria [40,41,42]. Key decomposed products of these reactions include phenols, amino acids, and fatty acids [31,43,44]. When thermal maturities greater than Ro~0.5% are reached, hydrocarbon molecules, organic acids, H₂O, CO₂, and N₂ within OM are released and have been described extensively as a mechanism for acidizing diagenetic systems [10,45,46,47,48]. Additionally, the dehydration of minerals such as clays and gypsum also affects the properties of diagenetic fluids [49,50,51]. Therefore, the complex diagenetic environment driven by these varying fluids highlights the inherent difficulties in resolving the complex origins of authigenic minerals in organic-rich shale successions [52,53], and these minerals are considered of significant importance for the reservoir quality of shales in many published studies [17,28,54]. Previous studies have mainly focused on authigenic minerals within marine shale, especially biogenic microquartz crystals, which generate abundant intercrystalline pores and are considered as critical factors enabling efficient shale gas production in both academia and industry [15,17,24,47]. However, related studies reported in lacustrine shales are limited. The complex petrology and fine grain size of lacustrine shales enhance the reactivity of their components, resulting in unique morphologies and distributions of authigenic minerals. These characteristics may lead to implications in reservoirs that are significantly different from those observed in marine shales.

The upper sub-member of the fourth member of the Eocene Shahejie Formation (Es₄^U) in the Dongying Sag of Bohai Bay Basin, China, is a typical saline lacustrine organic-rich shale succession and an important unconventional target [55,56]. This interval offers an excellent sample set to investigate the origin and diagenetic evolution of authigenic minerals and their impact on reservoir quality in lacustrine shales. This study focuses on detailed observations of core samples, thin sections, and scanning electron microscope (SEM + EDS), as well as quantitative analyses including X-ray diffraction (XRD), major and trace element analysis, total organic carbon (TOC) measurement, rock pyrolysis, and porosity analysis. The aim is to (i) decipher the formation mechanisms of authigenic minerals, (ii) reconstruct their evolutionary sequence, and (iii) assess their impact on lacustrine shale reservoirs. The findings are critical for reservoir quality evaluation and development optimization.

2. Geological Setting

The Dongying Depression is in the southeastern part of the Bohai Bay Basin, located in eastern China (Figure 1A), which is a typical half-graben rift basin (Figure 1B). The Dongying depression can be subdivided into seven secondary structural units, including the Minfeng Subsag, the Lijin Subsag, the Niuzhuang Subsag, the Boxing Subsag, the Southern Gentle Slope Zone, the Central Anticline Zone, and the Northern Steep Slope Zone, with an area of approximately 5700 km² (Figure 1C) [57]. The Paleogene strata within the study area are widely distributed and consist, from the bottom up, of the Kongdian Formation, the Shahejie Formation, and the Dongying Formation. The Shahejie Formation can be further divided into the fourth member (Es₄), the third member (Es₃), the second member (Es₂), and the first member (Es₁) [58,59]. The Es₄ is further divided into the Es₄^U and Es₄^L (Figure 1D). The Es₄^L primarily consists of mudstone, gypsum, coarse sandstone, and carbonate rock, while the Es₄^U is mainly composed of calcareous shale and dolomitic shale [58].

According to previous research, the Dongying Depression quickly subsided with a large-scale lake expansion, reaching a maximum during the Es₄^U. The saline, lacustrine, black organic-rich shale succession formed at the lake bottom (~50 m), which are the primary source rocks for the petroliferous depression [60,61]. The water salinity varied from semi-saline to saline within our target layer [62].

3. Samples and Experimental Methods

In total, 136 samples from the NX124 well (3213 to 3243.15 m), FY1 well (3248 to 3430 m), F120 well (3269.5 to 3289 m), and F143 well (3110 to 3141.08 m) were selected to analyze the genesis of authigenic minerals in organic-rich shale. The methods used included X-ray diffraction (XRD), total organic carbon content (TOC) measurement, rock pyrolysis, vitrinite reflectance (Ro) analysis, major and trace element analysis, sporopollen analysis, scanning electron microscopy (SEM + EDS), and helium porosity testing. These samples were obtained from China’s SINOPEC Shengli Oilfield Company and divided into bulk and cylindrical samples. The experiments were conducted in the Reservoir Geology and Basin Analysis Key Laboratory of the China University of Petroleum.

3.1. Petrographic Analysis

We selected 94 bulk samples from the FY1 well and 12 samples from the NX124 well for bulk-rock XRD analysis. First, we ground to powders of less than 300 mesh in size. The X’Pert PRO MPD with CuK α radiation (60 kV, 55 mA) and Ni filtering were used for X-ray diffraction (XRD) analyses to identify the bulk mineralogical composition. X-ray diffraction analysis was performed over a 2θ range of 1–160°, with mineral identification and semi-quantitative analysis based on characteristic peak intensities and positions. Subsequently, 43 samples from the FY1 well were separated for clay XRD. The samples were prepared as oriented mounts and subjected to ethylene glycol vapor treatment under 40–50 °C conditions without employing zincite (or similar) internal standards for quantitative calibration.

Thirty samples from the FY1 well were adhered to a standard glass slide with epoxy resin to analyze mineral and sporopollen characterization using a Leica DM2700P high-precision microscope (Leica Microsystems, Wetzlar, Germany).

3.2. Organic Geochemistry Analysis

The total organic carbon (TOC) contents of 102 bulk samples from the FY1 well and NX124 well were ground to powders of less than 200 mesh in size and treated with 12.5% hydrochloric acid to remove inorganic carbon. After rinsing with deionized water, the samples were filtered to achieve neutrality and dried in an oven at 60–80 °C. Finally, samples were analyzed using a Leco CS-744 carbon-sulfur analyzer with high-temperature oxygen flow combustion.

Rock pyrolysis was performed on 102 bulk samples from the FY1 well and NX124 well. The samples were ground to powders that were 200 mesh in size and placed in the Rock-Eval 6 analyzer. Then, samples were heated isothermally at 300 °C for three minutes to measure the free hydrocarbon content (S₁) released from the rock. Following this, the temperature was increased at a rate of 25 °C/min to 600 °C to obtain the pyrolyzed hydrocarbon content (S₂) generated from the thermal maturity of organic matter and to determine the temperature at which the maximum S₂ yield occurs (Tmax). The hydrogen index (HI) was calculated using the ratio of S₂ to TOC. The mixture was heated on a low-temperature hot plate at 140 °C for 48 h.

In addition, 22 samples from the FY1 well were separated for vitrinite reflectance to determine the diagenetic stage.

3.3. Inorganic Geochemistry Analysis

For major and trace element analysis, 12 samples from the NX124 well were ground to approximately 200 mesh. Then, about a 0.05 g sample was placed into a PTFE crucible with 25 mL of water, followed by the addition of 1 mL hydrochloric acid, 1 mL nitric acid, 2 mL hydrofluoric acid, and 0.5 mL perchloric acid. After evaporation to near dryness, 2 mL of a 1:1 nitric acid solution was added, and the sample was transferred to a 50 mL volumetric flask. The sample was then analyzed using an ELEMENT XR high-resolution inductively coupled plasma mass spectrometer (ICP-MS) (Thermo Fisher Scientific, Waltham, MA, USA).

3.4. SEM and EDS

The scanning electron microscope (SEM) and energy dispersive spectroscopy (EDS) analyses were performed on 12 samples selected from the NX124 well. The samples were first cut into slices with a diameter of approximately 1–3 cm and a thickness of about 0.2–1 cm. Then, 12 samples underwent mechanical and Ar-ion polishing to generate super-flat surfaces suitable for SEM examination. Subsequently, a thin layer of carbon film (approximately a few nanometers thick) was uniformly deposited on the sample surfaces. The Zeiss Crossbeam 550 FIB-SEM analyzer equipped with an energy dispersive spectrometer (EDS) was used as the observation instrument. In addition, 30 SEM samples from the FY1, F120, and F143 wells were analyzed, including Ar-polished images and fresh fracture surfaces images.

3.5. Helium Porosity

The helium porosity of 21 samples from the FY1 well and NX124 well was determined using the KXD-III helium porosity determinator to measure the porosity. Samples were prepared as cylinders with a diameter of 2.5 cm and height of 3 cm. To make helium fully fill the pores in the shale, the working pressure and equilibration time of the helium pycnometry test were 4 MPa for 8 min, respectively. Porosity was obtained from the difference between the bulk density and the skeletal density with a helium expansion method using Boyle’s Law helium porosimetry [44].

4. Results

4.1. Petrologic Features and Lithofacies Types

The mineral composition of shale, as shown in Table S1, is diverse in the study area (Figure 2A). The calcite content is the highest, ranging from 1% to 80%, with an average of 39.95%. Quartz content varies from 4% to 55%, averaging 24.52%. Clay minerals range from 3% to 54%, with an average of 16.85%, primarily consisting of illite (average of 86.81%), followed by mixed layers of illite-smectite (average of 12.95%) (Figure 2B). Dolomite content ranges from 1% to 93%, with an average of 11.65%. The contents of plagioclase and pyrite are relatively low, averaging below 5%. The sporopollen identification shows that its color is light yellow to black.

The Es₄^U is primarily characterized by calcareous shale, mixed shale, and dolomitic mudstone. Calcareous shale and mixed shale predominantly exhibit a well-lamellar texture, with thin sections revealing three-part laminations of bright carbonate laminae, brown felsic laminae, and dark organo-clay laminae. In contrast, dolomitic mudstone primarily exhibits a massive texture (Figure 2C).

4.2. Geochemical Characteristics

4.2.1. Organic Geochemistry

Organic geochemical data indicate that the TOC values of samples from Es₄^U range from 0.55% to 9.05%, with an average of 2.27% (Table S2). Among these, 99.02% of the samples have values greater than 1%, and 51.96% have values greater than 2%. The S₁ values range from 0.34 to 5.8 mg/g, with an average of 2.28 mg/g (Figure 3A), while S₂ values range from 0.93 to 41.08 mg/g, averaging 7.49 mg/g (Figure 3B). The oil saturation index (OSI) can be used to characterize the mobility efficiency of retained hydrocarbons in shale reservoirs [63]. When the OSI exceeds 100, hydrocarbons exceeding the adsorption capacity of the organic matter will exhibit better mobility and oil production potential [19,64]. The OSI values of samples range from 28.82 to 377.52 mg/g, with an average of 107.86 mg/g, indicating that the Es₄^U shale has good oil potential. The Tmax values range from 431 to 461 °C, with an average of 448.04 °C. The HI values range from 129.17 to 544.69 mg/g, averaging 319.66 mg/g. According to the cross plot of Tmax and HI (Figure 3C), the OM is primarily comprised of type I and minor type II. The measured vitrinite reflectance (Ro) values of 22 samples range from 0.52% to 0.93%, indicating a thermal maturity stage spanning from low maturity to the oil window (Table S3).

4.2.2. Elemental Geochemistry

The major element of samples is shown in Table S4. The most abundant element is SiO₂, with a content ranging from 27.4% to 50.2% and an average of 37.76%. The next most abundant elements are CaO and Al₂O₃, with average contents of 18.26% and 10.43%, respectively. The contents of Fe₂O₃, MgO, and K₂O range from 1% to 5%, while Na₂O is approximately 1%. The contents of TiO₂, P₂O₅, and MnO are generally below 1%.

The trace elements of samples are shown in Table S5. The most abundant trace elements are Sr and Ba, with average concentrations of 1985.58 ppm and 692.25 ppm, respectively. Following these are Li, B, V, Cr, Co, Ni, Cu, Zn, Ga, Rb, Y, and Zr, which have concentrations ranging from 10 to 100 ppm. Be, Sc, As, Nb, Mo, Cd, and Cs have the lowest concentrations, generally falling below 10 ppm.

4.3. Types and Characteristics of Authigenic Minerals

4.3.1. Characteristics of Authigenic Quartz

Authigenic quartz widely appears in forms with short columnar and hexagonal columnar clay-size euhedral microquartz, primarily developing within intergranular pores and adjacent to calcite and illite (Figure 4A).

4.3.2. Characteristics of Pyrite

Two distinct pyrite morphologies dominate the shale samples. Framboidal aggregates are composed of closely packed, submicron equigranular crystals, forming elliptical clusters (Figure 4B). Euhedral pyrites often appear as discrete crystals with cubic and occasionally elliptical morphologies, with the grain sizes ranging from 1 to 5 μm (Figure 4C). The pyrite concentrations in shales range from 0% to 34%, with an average of 2.84%.

4.3.3. Characteristics of Ferroan Dolomite

The ferroan dolomite is in the form of micron-sized euhedral crystals, which are comprised of an irregular core of dolomite and an annular rim of ankerite. Ferroan dolomite is commonly observed within clays, with grain size ranging from 2 to 20 μm (Figure 4D). The crystal surfaces are flat and not curved, exhibiting a relatively complete rhombohedral shape.

4.3.4. Characteristics of Authigenic Clay Minerals

The authigenic clays are primarily comprised of kaolinite and chlorite. The kaolinite appears in book-like and leaf-like shapes, often accompanied by microquartz and dissolution of K-feldspar (Figure 4E,F). Chlorite usually appears in hairy and fibrous filaments (Figure 4G,H). The content of kaolinite and chlorite is limited, occurring only at specific depths (Figure 2B).

4.3.5. Characteristics of Albite

The albite is primarily in the form of small granular and columnar euhedral crystals, with grain sizes ranging from 5 to 10 μm. The albite is commonly observed filling the intergranular pore in association with dissolved K-feldspar and illite (Figure 4I). Compared to other feldspar, the albite was characterized by weak dissolution with no evidence of replacement or cementation by other minerals.

4.4. Porosity Characteristics

Twenty-one samples revealed the porosity characteristics of the Es₄^U shale, as shown in Table S1. The porosity varies greatly in different wells. Among them, nine samples from the FY1 well showed porosity ranging from 2.65 to 9.43%, with an average of 6.13%. Twelve samples from the NX124 well showed porosity ranging from 7.37 to 15.6%, with an average of 11.25%. Both of them can be classified as low porosity reservoirs.

4.5. Diagenetic Stage

This study used Ro, Tmax, clay mineral types and assemblages, authigenic mineral types, pore types, and sporopollen color as primary indices for classifying the diagenetic stage [65,66,67]. The Ro ranges from 0.52% to 0.93%, and Tmax ranges from 415 to 451 °C, indicating that the maturity of shales is in the low-mature to mature stage. The clay minerals are composed of illite and illite-smectite mixed layers, with the ratio of the illite-smectite mixed layer being 20%, suggesting an ordered mixed layer. The sparry calcite and ferroan dolomite with bright rims are widely observed, and pore types are mainly classified as dissolved pores and recrystallized intergranular pores in samples. The sporopollen color ranges from light yellow to black. Therefore, we conclude that the Es₄^U shale corresponds to the mesodiagenetic A stage (

Table 1. Diagenetic division indices of the Es₄^U shale.

Dividing Index	Number	Distribution Range (or Main Type)	Diagenetic Stage
Ro/%	22	0.52–0.93	Mesodiagegesis A
Tmax/°C	90	431–461 °C	Mesodiagegesis A
Type and combination of clay minerals	43	(I/S + I + C)	Mesodiagegesis A
	43	The ratio of I/S is 20%	Mesodiagegesis A
	43	Ordered I/S	Mesodiagegesis A
	30	Acicular and flocculent	Mesodiagegesis A
Type and crystallinity of carbonate minerals	30	Sparry calcite	Mesodiagegesis A
	30	Ferroan dolomite	Mesodiagegesis A
Type of pores	30	Dissolved pores and recrystallized intergranular pores	Mesodiagegesis A
Color of sporopollen	39	Light yellow to black	Mesodiagegesis A

).

5. Discussion

5.1. The Origin of Authigenic Minerals

5.1.1. The Origin of Authigenic Quartz

The extensively exposed continuous deep-water mudstone sequences lack definitive petrological evidence for active hydrothermal or volcanic activity in the study area [60,61]. Hydrothermal authigenic quartz typically occurs as coarse-grained anhedral crystals or vein fillings [68]. However, SEM imaging reveals euhedral, clay-sized microquartz crystals spatially associated with clay minerals (e.g., kaolinite, illite) and dissolved K-feldspar (Figure 5A,B), indicating silica release linked to smectite illitization and acidic dissolution of K-feldspar. Additionally, quartz grains adjacent to pyrite and illite exhibit dissolution features (irregular shapes, diffuse boundaries) (Figure 5C,D) consistent with alkaline-mediated dissolution of detrital quartz precursor [69]. The study area represents a saline lacustrine basin with strongly alkaline conditions (Figure 6), where detrital quartz grains are susceptible to dissolution.

Following the principle of mass balance, various silicon components can be quantitatively determined using major element analysis combined with XRD data [70,71]. The calculation of detrital silica (Si_det) is given by the formula Si_det = Al_sample ∗ (Si/Al)_background, where (Si/Al)_background is the ratio between Si and Al in the background detrital flux, which is approximately 3.11 [72]. Additionally, a shale containing 787.2 g illite can generate 197.7 g SiO₂ during the smectite illitization [70]. Therefore, the silica released from the smectite illitization (Si_clay) can be calculated. Finally, the remaining excess silica (Si_excess) can be calculated using the formula (Si/Al)_total − 3.11/(Si/Al)_total [71]. The results are shown in

Table 2. Calculated percentages of quartz of different origins of shales from Es₄^U.

Samples	Depth/m	Al/%	Si/%	(Si/Al)total	Sidet/%	Siclay/%	Siexcess/%	Sitotal/%
NX124-2	3213.00	10.70	41.00	3.83	33.28	4.56	0.19	38.02
NX124-4	3214.15	9.86	34.80	3.53	30.66	4.69	0.12	35.48
NX124-7	3216.00	11.10	38.40	3.46	34.52	5.63	0.10	40.25
NX124-10	3218.00	7.54	27.40	3.63	23.45	2.89	0.14	26.49
NX124-15	3221.00	12.50	42.70	3.42	38.88	5.93	0.09	44.90
NX124-19	3223.00	11.90	43.50	3.66	37.01	5.15	0.15	42.31
NX124-23	3226.25	10.70	36.30	3.39	33.28	5.22	0.08	38.58
NX124-28	3229.10	8.44	34.00	4.03	26.25	4.31	0.23	30.79
NX124-32	3231.60	6.90	34.50	5.00	21.46	3.83	0.38	25.67
NX124-38	3234.15	8.24	32.70	3.97	25.63	3.71	0.22	29.55
NX124-47	3238.70	11.20	37.60	3.36	34.83	4.70	0.07	39.61
NX124-57	3243.15	16.10	50.20	3.12	50.07	10.26	0.00	60.33

. The content of Si_det ranges from 21.46% to 50.07%, Si_clay ranges from 2.89 to 10.26, and Si_excess ranges from 0 to 0.38. The calculated Si_total shows a good correlation with measured Si, and their values are relatively close (Figure 7A). The lacustrine shales in our study area show no direct evidence (e.g., sponge spicules or radiolarians) of biogenic silica. However, given documented occurrences of silica fauna in lacustrine shale [73], a biogenic silica component cannot be entirely excluded. The SiO₂-Zr correlation serves as a reliable proxy for silica provenance, where positive correlation reflects terrigenous input and negative correlation implies biogenic sources [73]. Our geochemical data reveal strong Zr correlations with Si_det and Si_clay (confirming terrigenous derivation), contrasted by a clear negative Zr-Si_excess relationship (pointing to potential biogenic contributions) (Figure 7B). Biogenic silica, if present at all, constitutes a volumetrically negligible component (<0.1%) of the total authigenic quartz inventory, with terrigenous sources and diagenetic transformations dominating the silica inventory.

5.1.2. The Origin of Pyrite

Pyrites predominantly appear in samples as framboidal and euhedral pyrites, which are often observed in other marine and lacustrine organic-rich shale successions [14,28,72,74]. According to the cross plot of grain size distribution and standard deviation, pyrites in samples have a grain size of less than 5 μm (Figure 8), primarily forming in a sulfurized environment [75]. The supply of S was derived from bacterial sulfate reduction (MSR) with shallow burial and thermochemical sulfate reduction (TSR) with deep burial. These processes were controlled by the openness of the diagenetic system, solute supply, and temperature and pressure conditions [76,77]. Framboidal pyrites primarily form during the syngenetic or early diagenetic stages [78], which have a strong sulfidic environment and a relatively open diagenetic system. Abundant Fe²⁺ and HS⁻ allow pyrite microparticles to rapidly nucleate, grow, and aggregate, forming an equigranular structure [74,75,79]. Euhedral pyrites form in two stages: (i) directly precipitated from unsaturated solutions of Fe and S during the syngenetic stage [80] and (ii) crystal aggregation and growth of framboidal pyrites during the early stage [38,81]. These origins emphasize that the solute supply in unsaturated fluids is limited, primarily formed by crystal overgrowth [82]. A third formation pathway involves pyrite generation via thermochemical sulfate reduction (TSR) after the mesodiagenetic stage [83]. However, TSR requires higher temperatures and typically occurs after mesodiagenesis [77,83]. The shale of Es₄^U is in the mesodiagenetic A stage, which may not exhibit the temperature and pressure conditions required for this reaction.

5.1.3. The Origin of Ferroan Dolomite

Dolomite is difficult to directly form at normal temperatures. Previous studies indicate that the minimum temperature required for dolomite precipitation in the laboratory is approximately 50–60 °C [84,85]. However, microbial mediation can overcome limitations that dolomite can directly precipitate at room temperature [86,87]. Analogous studies have been reported in the global Precambrian sedimentary rock record [88,89,90,91]. Ferroan dolomite is generally considered a product of hydrothermalism and burial dolomitization [92,93]. Hydrothermal ferroan dolomite exhibits coarse grain with a saddle texture and an obvious bright rim, with higher Fe contents usually exceeding 50% [94,95,96]. However, observations in samples indicate that ferroan dolomite crystals exhibit fine grains with indistinct bright rims and Fe content that is lower than 15% (Figure 9); thus, it is unlikely to be of hydrothermal origin. Therefore, it is more likely that the ferroan dolomite in samples is primarily a product of burial dolomitization. The Mg²⁺ and Fe²⁺ released from the smectite illitization may serve as the major source of ferroan dolomite [97]:

$$4.5{\mathrm{K}}^{+}+8{\mathrm{Al}}^{3+}+\mathrm{Smectite}\to \mathrm{Illite}+{\mathrm{Na}}^{+}+2{\mathrm{Ca}}^{2+}+2.5{\mathrm{Fe}}^{3+}+2{\mathrm{Mg}}^{2+}+3{\mathrm{Si}}^{4+}$$

(1)

$$\mathrm{Calcite}+{\mathrm{Fe}}^{2+}+{\mathrm{Mg}}^{2+}\to \mathrm{Ferroan} \mathrm{dolomite}$$

(2)

The cross plot of the dolomite content versus values of Mg + Fe reveals a significant positive correlation (Figure 10A), indicating that dolomite is in the main form of ferroan dolomite. There is also a positive correlation between Mg²⁺ and Fe²⁺ (Figure 10B), suggesting that their sources and enrichment characteristics are likely similar. Additionally, the positive correlation between K and the values of Fe + Mg also indicates a related origin of these ions (Figure 10C). The K-rich diagenetic system is in favor of the smectite illitization that promotes the release of Fe²⁺ and Mg²⁺ as well as the precipitation of ferroan dolomite [98,99]. This is obviously observed in samples where ferroan dolomite is widely distributed within clays.

Additionally, the dissolution of calcite and chlorite driven by CO₂ can induce precipitation of ferroan dolomite:

$$\mathrm{Chlorite}+5{\mathrm{CaCO}}_3+5{\mathrm{CO}}_2\to \mathrm{Ferroan} \mathrm{dolomite}+\mathrm{Kaolinite}+{\mathrm{SiO}}_2+2{\mathrm{H}}_2\mathrm{O}$$

(3)

In relatively closed diagenetic systems of shale, CO₂ is related to hydrocarbon generation [100]. The dissolution of chlorite releases Mg²⁺ and Fe²⁺, contributing to the precipitation of ferroan dolomite [101]. Stronger dissolution of chlorite results in increased ferroan dolomite [102], with the reaction being influenced by temperature and pH. The optimal temperature range for this reaction is between 100 °C and 140 °C, and suitable pH conditions are neutral to acidic [103]. Burial temperature can be calculated using vitrinite reflectance (Ro):

$$\mathrm{T}=\left(\mathrm{In}\right(\mathrm{Ro}\%)+1.26)/0.0081$$

(4)

The values of Ro range from 0.52% to 0.93%, corresponding to calculated temperatures between 120.49 °C and 151.66 °C, which partially fall within the optimal temperature range for this reaction. The cross plot of the calcite and ferroan dolomite reveals an obvious negative correlation (Figure 10D), indicating that the above process probably occurs in diagenetic history. However, the dissolution of chlorite is a slow process [104], and its content is relatively low (Figure 2B). The amount of ferroan dolomite formed through this process may be limited.

5.1.4. The Origin of Clay Minerals

Clay minerals in sedimentary basins primarily originate from terrestrial input and diagenetic alteration [98,105,106]. Illite-smectite mixed layers and illite have extremely high content (Figure 2B), closely related to alkaline connate waters [12,107,108]. Smectite rapidly transforms into illite with burial. Therefore, the precursor of current illite-smectite mixed layers and illite is early smectite. Kaolinite is a product of alteration of K-feldspar grains, which originate from thermodynamic reactions of feldspar dissolution and reprecipitation and are stable in an acidic diagenetic system [98,109].

$$2\mathrm{K-feldspar}+2{\mathrm{H}}^{+}+{\mathrm{H}}_2\mathrm{O}\to \mathrm{Kaolinite}+4{\mathrm{SiO}}_2+2{\mathrm{K}}^{+}$$

(5)

$$2\mathrm{Albite}+2{\mathrm{H}}^{+}+{\mathrm{H}}_2\mathrm{O}\to \mathrm{Kaolinite}+4{\mathrm{SiO}}_2+2{\mathrm{Na}}^{+}$$

(6)

$$\mathrm{Anorthite}+2{\mathrm{H}}^{+}+{\mathrm{H}}_2\mathrm{O}\to \mathrm{Kaolinite}+{\mathrm{Ca}}^{2+}$$

(7)

Photomicrographs revealed that kaolinite exhibits a booklet shape that grows with microquartz, and obvious dissolved K-feldspar with microquartz was observed in the sample (Figure 5B). The cross plot of kaolinite and feldspar reveals a negative correlation (Figure 11A), further suggesting that the development of kaolinite is primarily related to the acidic dissolution of K-feldspar.

Chlorite fibers are primarily products of kaolinite chloritization. Diagenetic fluids change and enrich Mg²⁺ through dissolved dolomite and smectite illitization in mesodiagenesis, triggering the kaolinite chloritization [110]. The cross plot of kaolinite and chlorite indicates a negative correlation (Figure 11B), suggesting the growth of chlorite as dissolution of kaolinite. Photomicrographs reveal that chlorite exhibits hairy and fibrous crystals (Figure 4G,H), indicating its diagenetic origin.

5.1.5. The Origin of Albite

Albite in shales has various origins. The occurrence of volcanic and hydrothermal albite in organic-rich lacustrine shales from the Permian Lucaogou Formation has been well documented [68,111]. However, there is not an obvious volcanic and hydrothermal filling record in Es₄^U shale. SEM observations indicate that albites exhibit well-euhedral crystals with fine-grain size, widely occurring with calcite and illite (Figure 4I). These are interpreted as alkaline alteration products of feldspar [25,69].

$$\mathrm{Anorthite}+2{\mathrm{Na}}^{+}+4{\mathrm{SiO}}_2=2{\mathrm{NaAlSi}}_3{\mathrm{O}}_8+{\mathrm{Ca}}^{2+}$$

(8)

$$2\mathrm{K-feldspar}+2.5\mathrm{kaolinite}+2{\mathrm{Na}}^{+}+4{\mathrm{SiO}}_2=2\mathrm{albite}+2\mathrm{illite} +5{\mathrm{H}}_2\mathrm{O}+2{\mathrm{H}}^{+}$$

(9)

The alkaline alteration of K-feldspar and anorthite requires a sufficient supply of SiO₂ and Na⁺. This process works synergistically with the reaction in Equation (1), where the release of Si⁴⁺ and Na⁺ from smectite illitization promotes the alkaline alteration of K-feldspar and anorthite. The cross plot of Na⁺ versus Fe²⁺ and Mg²⁺ indicates a positive correlation (Figure 12A), suggesting that Na, Fe, and Mg may all originate from the smectite illitization. The content of K-feldspar is very low and predominantly features plagioclase in shales (Table S1), which have a strong positive correlation with Na⁺ (Figure 12B). SEM observations reveal that albite is the only observable plagioclase phase in samples (Figure 13). However, feldspar from terrestrial input is unlikely to consist entirely of albite. It is concluded that albite primarily originates from the alkaline alteration of K-feldspar and anorthite.

5.2. The Diagenetic Evolution Sequence of Authigenic Minerals

Organic-rich shale has undergone a lengthy and complex process of organic and inorganic diagenesis, which is in the stage of mesodiagenesis A (Table 1). Based on the morphology, mineral assemblage, and diagenetic features of authigenic minerals, their diagenetic evolution sequence has been established (Figure 14).

During the syngenetic diagenesis stage, clay-rich sediments are rich in connate sedimentary water, with primary porosity reaching up to 70% [112], indicating an open diagenetic system. The diagenetic fluid consists of inherited alkaline primary sedimentary water. Due to the abundance of solutes, particularly S and Fe, framboidal pyrite preferentially precipitates [74,79]. The framboidal pyrite crystals were uniform in size and distributed closely, indicating a relatively open diagenetic system.

In the eodiagenetic A stage, the paleotemperature is less than 65 °C, and Ro is below 0.35%. Calcareous sediments rapidly dewater under the compaction, leading to the closure of the diagenetic system. As the abundance of S and Fe solutes decreases, the formation of framboidal pyrite becomes difficult, and precipitation in the form of euhedral pyrite is noted [32]. The euhedral pyrite crystals were of nonuniform sizes, and their distribution was controlled by pores, indicating a relatively closed diagenetic system. Felsic grains dissolve in an alkaline environment, with ions such as Si⁴⁺, K⁺, Na⁺, and Ca²⁺ transferring into the diagenetic fluids.

In the eodiagenetic B stage, the paleotemperature ranges from 65 to 85 °C, and Ro ranges from 0.35 to 0.5%. Here, mechanical compaction is the primary factor leading to a closed diagenetic system [4,11]. More ions are released by smectite illitization, such as Si⁴⁺, Mg²⁺, Na⁺, Ca²⁺, and Fe²⁺, into diagenetic fluids [98,113]. The synergistic effect of smectite illitization that continuously consumes K⁺, as well as alkaline alteration of feldspar that continuously consumes Na⁺, leads to a significant decrease in K-feldspar and anorthite, with an increasing number of albites. Ca²⁺, Mg²⁺, and Fe²⁺ of diagenetic fluids will precipitate as ferroan dolomite. When the temperature was about 70 °C, a few organic acids were released by the hydrocarbon generation of organic matter, which acidized the diagenetic environment [45,114]. Any silica precipitated from diagenetic fluids typically forms as microquartz.

In the mesodiagenetic A stage, the paleotemperature ranges from 85 to 140 °C, and Ro ranges from 0.5 to 1.3%. The organic acids and CO₂ released by organic matter acidize the diagenetic system, which dissolves feldspar and carbonate minerals [115,116]. Dissolved K-feldspar was accompanied by microquartz and kaolinite precipitate in pores, while the release of K⁺ promotes smectite illitization [46,117] and further precipitates as ferroan dolomite [93,99]. Additionally, the Mg²⁺ enriched in diagenetic fluid results from smectite illitization and the dissolution of dolomite, resulting in kaolinite chloritization [103,110]. Small amounts of chlorite and calcite dissolve and precipitate as kaolinite and ferroan dolomite [102,104].

5.3. Implications for Shale Reservoir Quality

Authigenic minerals in shales exhibit complex organic and inorganic interactions that critically influence reservoir quality. The cross plot of diagenetic silica, ferroan dolomite, albite, and pyrite versus porosity shows a relatively positive correlation (Figure 15A–D), suggesting that these minerals have significant implications for shale reservoirs. Although they filled most of the pores, they exhibit important clues for understanding the mechanisms that preserve pores during burial. Beyond overpressure, grain support frameworks constitute the dominant mechanism preserving open intergranular pores throughout shale diagenesis [10,14,118]. Unlike in marine shales where authigenic quartz contributes to pore development [24,38,119], lacustrine authigenic quartz primarily enhances reservoir quality through improved brittleness rather than porosity creation (Figure 4A). While framboidal pyrite develops intercrystalline pores that enhance porosity (Figure 4B), euhedral pyrite promotes hydrocarbon cracking through surface iron ions despite its inability to form pores [100,120]. In some well-documented studies, dolomite, rather than calcite, dominates dissolution within shale reservoirs [118,121]. However, the specific mechanisms are not yet comprehensively understood. A possible explanation was that Mg and Fe incorporation changed the free energy within the crystal and enhanced mineral solubility [122,123,124]. Thus, the dissolution of dolomite may generate more porosity rather than the dissolution of calcite during burial, although dolomite is less common than calcite. Theoretically, alkaline alteration of K-feldspar and anorthite can generate porosity values of 15.4% and 6.92%, respectively [125]. It looks favorable, but the actual content of feldspar is low. Calcites, dolomites, and microquartz crystals are the dominant non-clay minerals in organic-rich shales [34,63,126]. Feldspar grains are chemically unstable and undergo dissolution, forming dissolved pores during deep burial [127]. Based on the photomicrographs, however, dissolved pores are not well-developed in albite (Figure 4I and Figure 14). This process is associated with an increased concentration of Na⁺ within diagenetic fluids, which in turn leads to albite precipitation but not dissolution [128]. Nevertheless, albites still have the potential to enhance the mechanical strength of rock and preserve porosity. As products of diagenetic alteration, authigenic clay minerals degrade pore structure and intensify pore throat heterogeneity [129,130] (Figure 15E,F), but they may contribute to hydrocarbon adsorption and enrichment [48,131].

This study seeks to decode the origin and diagenetic evolution of authigenic minerals in lacustrine shale to evaluate their implications on reservoir quality. However, the number of samples (12) is insufficient to accurately illustrate the complexity and heterogeneity of authigenic minerals in microscale diagenetic systems. Moreover, the methods used are rather conventional. Considerably more sample sets and high-resolution tests are required for a comprehensive understanding of the multiple origins of authigenic minerals. In particular, in-situ high-resolution crystallographic, mineralogical, and element geochemical data are needed to allow for more precise correlations of nucleation, growth, and dissolution of authigenic minerals across the diagenetic history.

6. Conclusions

This study systematically investigates the formation mechanisms of authigenic minerals in lacustrine organic-rich shales and their multiscale impacts on reservoir quality.

1. Authigenic minerals in lacustrine organic-rich shale are dominated by microquartz, framboidal and euhedral pyrites, ferroan dolomite, kaolinite, chlorite, and albite. These authigenic minerals mostly display submicron-sized euhedral crystals and are distributed in large pores.

2. Euxinic and saline primary sedimentary water dominated the type and abundance of early diagenetic authigenic minerals. Framboidal pyrites precipitated directly from solutions with high sulfur saturation. When the sulfur saturation decreased, framboidal pyrites could further grow and transform into euhedral pyrites. Persistent alkaline alteration of feldspar could be an important mechanism for albite growth and preservation.

3. Acid fluids released by the maturity of organic matter are responsible for the formation of an acidified diagenetic system. The dissolution of K-feldspar and carbonate minerals, as well as smectite illitization, provided ions for a variety of mesodiagenetic mineral cements, which resulted for instance in microquartz crystals, ferroan dolomites, kaolinite, and chlorite cemented in pores.

4. Although most rigid authigenic minerals directly precipitate in intergranular pores, they most likely contributed to reservoir quality via a slight reduction in the impact of compaction.

5. This study can be used to investigate authigenic minerals in comparable lacustrine organic-rich shale. However, significantly larger sample sets and in-situ high-resolution tests would be required to conduct case studies and assess their influence on shale reservoirs.