Recognition of diagenetic contribution to the formation of limestone-marl alternations: A case study from Permian of South China

Marine and Petroleum Geology 111 (2020) 765–785

Contents lists available at ScienceDirect

Marine and Petroleum Geology journal homepage: www.elsevier.com/locate/marpetgeo

Research paper

Recognition of diagenetic contribution to the formation of limestone-marl alternations: A case study from Permian of South China

T

Chengpeng Sua,b, Fei Lia,c,∗, Xiucheng Tana,c,∗∗, Qiaolin Gongb, Kai Zengb, Hao Tanga,c, Minglong Lib, Xiaofang Wangb,d a

State Key Laboratory of Oil and Gas Reservoir Geology and Exploitation, Southwest Petroleum University, Chengdu, 610500, China School of Geoscience and Technology, Southwest Petroleum University, Chengdu, 610500, China c Department of Key Laboratory of Carbonate Reservoirs of CNPC, Southwest Petroleum University, Chengdu, 610500, China d Key Laboratory of Carbonate Reservoirs, CNPC, Hangzhou, 310023, China b

ARTICLE INFO

ABSTRACT

Keywords: Limestone-marl rhythmites Carbonate diagenesis Siliciclastic contamination REE+Y LA-ICP-MS

Limestone-marl alternations (LMAs) have been extensively used in studies of sedimentology, cyclostratigraphy, and paleoclimatology. However, it has long been debated whether LMAs are derived from primary sedimentation or from diagenetic alteration, and the markers to distinguish one source from the other remain ambiguous. In this study, we investigate widespread Permian LMAs from various depositional settings in South China using traditional petrographic, mineralogical, and major element (Al2O3 vs.TiO2) methods, as well as emerging trace and rare earth element (REE) analyses performed using solution-based and in situ techniques. The original depositional differences of the limestone and its coupled marl beds can be identified based on (1) the high volumes of siliciclastic minerals in the marls, (2) unremarkable differential compaction, (3) the recognizable signatures of the detrital admixture (i.e., the Al, Zr, and Th characteristics), and (4) flattened shalenormalized REE patterns. Diagenetic LMAs are characterized by the compaction of fossils, rare aragonite skeletons, and abundant diagenetic clays (sepiolite and talc) preserved in the marl beds. We also identified fully diagenetic LMAs originating from primary carbonate sedimentation based on two almost completely consistent seawater-like REE patterns with diagnostic positive La and negative Ce anomalies ((La/La*)SN = 1.64 ± 0.17, (Ce/Ce*)SN = 0.73 ± 0.05), superchondritic Y/Ho (47 ± 4), and depleted light REEs relative to high REEs ((Pr/Yb)SN = 0.39 ± 0.09). Diagenesis also dominates the formation of LMAs with limited terrigenous input in shallow-water carbonate depositional environments. It is likely that both the active metabolic processes in early diagenesis (especially aerobic sulfide oxidation) and the temporal palaeoceanographic conditions, e.g., high levels of dissolved oxygen and sulfate, play crucial roles in the genesis of the diagenetic LMAs in the Permian strata of South China.

1. Introduction Limestone‒marl alternations (LMAs) refer to two stacked natural successions, which can be subdivided into alternating limestone and lime mudstone, limestone and argillaceous limestone, and limestone‒shale sedimentation based on the proportion of carbonate within the marl components (Einsele and Ricken, 1991; Munnecke and Samtleben, 1996; Westphal and Munnecke, 2003). Morphologically, LMAs are generally composed of well-to irregularly-bedded successions (Einsele and Ricken, 1991; Westphal and Munnecke, 2003; Westphal et al., 2010; Amberg et al., 2016). They can form in various settings (tidal flats to pelagic oceans; Einsele et al., 1991; Elrick and Hinnov, ∗

2007; Westphal et al., 2008a; Eldrett et al., 2015a). A systematic review of the LMAs literature reveals that their temporal distributions range from Cambrian to recent strata. They seem to be more concentrated in warm-waters and greenhouse conditions based on ancient records (Westphal and Munnecke, 2003; Westphal, 2006). Successive LMA deposits have been regarded as a result of orbitforced cyclic sedimentation (Fischer and Bottjer, 1991; Elder et al., 1994; Dinarès-Turell et al., 2003, 2010; Strasser et al., 2006; Berrocoso et al., 2013; Ma et al., 2017), and thus, they can provide useful information on the periodic oscillations observed in various fields of study, e.g., cyclostratigraphy (Hilgen, 1987; Montgomery et al., 2001; Roth and Reijmer, 2005; Eldrett et al., 2015b), climate (Gardulski et al.,

Corresponding author. #8 Xindu Road, Xindu District, Chengdu, 610500, China. Tel.: +8617358503252. Corresponding author. #8 Xindu Road, Xindu District, Chengdu, 610500, China. Tel.: +8613408653780. E-mail addresses: [email protected] (F. Li), [email protected] (X. Tan).

∗∗

https://doi.org/10.1016/j.marpetgeo.2019.08.033 Received 19 April 2019; Received in revised form 8 July 2019; Accepted 19 August 2019 Available online 22 August 2019 0264-8172/ © 2019 Elsevier Ltd. All rights reserved.

Marine and Petroleum Geology 111 (2020) 765–785

C. Su, et al.

1990; Bauer et al., 1996; Elrick and Hinnov, 2007; Bádenas et al., 2012; Lathuilière et al., 2015), carbonate productivity (Seibold, 1952; Bellanca et al., 1996, 1997), terrestrial siliciclastic input (Einsele and Seilacher, 1982; Bellanca et al., 1996; Eldrett et al., 2015a), and sealevel fluctuations (King, 1990; Pittet et al., 2000; Boulila et al., 2010; Coimbra et al., 2015). However, a substantial proportion of LMA successions, especially those from shallow marine settings, manifest as carbonate-dominated LMA sequences, e.g., alternating limestone and lime mudstone; and alternating limestone and argillaceous limestone. Prevalent carbonate diagenesis affects the mineralogical composition, fossil types and distributions, dissolution and precipitation, etc. Thus, it may play an essential role in the formation of dissolved and altered marl beds from original limestone deposits, which differs from the idea of an origin from physical compaction enforced by the overlying carbonate/siliciclastic components (Reinhardt et al., 2000; Westphal et al., 2000). In addition, the argillaceous components may participate to form secondary replacement products and authigenic minerals during early diagenesis, e.g., the transformation of clay minerals (Shaw and Primmer, 1991). Several studies argued that diagenetic alteration is inconspicuous (or minor) in the primary succession of LMA (Ricken, 1986; Bellanca et al., 1996; Huang and Baus, 1999; Husson et al., 2014; Eldrett et al., 2015a; Elderbak and Leckie, 2016), while others believe that the differential diagenesis occurring in the two neighboring beds would significantly affect relevant analyses based on the primary sedimentary composition, and thus, it needs to be carefully investigated (Hallam, 1986; Raiswell, 1987; Munnecke et al., 1997; Reinhardt et al., 2000; Holmes et al., 2004; Beltran et al., 2009; Westphal et al., 2010, 2015). The available methods of distinguishing between the original and diagenetic signatures of preserved LMAs remain controversial. Several lines of evidence from morphology, paleontology (e.g., palynomorphs), and major element geochemistry (e.g., Al2O3 vs. TiO2) have been used to determine the signals of early diagenesis in the formation of rhythmite (Hattin, 1971; Courtinat, 1993; Waterhouse, 1999; Westphal et al., 2000, 2004), but their applicability has recently been questioned (Westphal et al., 2010; Amberg et al., 2016). For example, palynomorph assemblages, which are regarded as diagenetically inert compounds, can be oxidized at high temperatures and compacted within the marls during the burial diagenetic phase (Amberg et al., 2016). In addition, the nearly uniform slope of Al2O3 vs. TiO2 within an LMA bed indicates either a primary or an overprinted diagenetic signal (Westphal et al., 2008a, 2010). It is possible to use the rare earth element and yttrium (REE + Y) compositions of the different LMA components to distinguish between the potential impacts triggered by primary depositional detritus or by carbonate diagenesis. The carbonate REE + Y distributions are characterized by very similar geochemical behaviors, except for Ce and Y anomalies, which may be related to the nature of the ambient water during precipitation (e.g., Li et al., 2019). The preservation of original REE signatures of carbonate components principally depends on the pore-water conditions and the dissolution and reprecipitation of carbonate minerals during diagenesis (Haley et al., 2004; Webb et al., 2009; Planavsky et al., 2010). Ancient shallow-water “pure carbonate” (defined as extremely low or almost negligible terrigenous input in this study) rocks with weak diagenetic overprints generally exhibit diagnostic heavy REE (HREE)-enriched pattern relative to light REEs (LREEs) when normalized to shales (Nothdurft et al., 2004). Moreover, carbonate REE compositions are sensitive to increases in siliciclastics, which lead to growing modifications on the shale-normalized REE patterns (especially for the LREEs), as well as smaller Ce (to 1) and Y anomalies (Kamber et al., 2004; Li et al., 2019). It is expected that the completely diagenetic process would not markedly modify the REE signatures if the marls originated from primary carbonate sedimentation. In contrast, the REE signatures of marls derived from mixed carbonate-siliciclastic components would be overprinted by terrigenous signatures and exhibit flattened LREE patterns (or enriched middle REE

(MREE) patterns), increased total REE (ΣREE) concentrations, and smaller Ce and Y anomalies depending on the specific diagenetic processes and the detrital volumes (Sholkovitz et al., 1994; Kamber et al., 2004; Zhao and Zheng, 2017; Li et al., 2019). To better understand the potential influences of carbonate diagenesis and siliciclastic input on the formation of LMAs, in this study, we used samples from different depositional settings (onshore, platform interior, and deep shelf), with different terrigenous supplies (rare to abundant), and different morphological types (well-bedded vs. nodularlike) from the Permian strata in South China. The aim of this study is to identify and distinguish the original diagenetic signatures of preserved LMA successions. To achieve this goal, traditional petrological, mineralogical, and major-element geochemical analyses were adopted to explore the different characteristics of the coupled lime and marl beds within the LMAs. In situ laser ablation–inductively coupled plasma–mass spectroscopy (LA-ICP-MS) and solution-based ICP-MS analyses were also used to determine the trace and rare earth element compositions of the lime and marl components, respectively. 2. Location and geological setting The South China Craton was located in the paleo-equatorial region during the Permian (~299−252 Ma) (Fig. 1) (Scotese and Langford, 1995). It consisted of two collided blocks (i.e., Yangtze and Cathaysia) and two deep-water basins (Jiangnan and Youjiang) (Wang et al., 1994). At that time, the Yangtze block was principally covered by extensive shallow-marine sedimentation. The block could be further subdivided into the upper (western), middle (central), and lower (eastern) portions based on their temporal geographic positions (Fig. 1B) (Feng et al., 1996; Wang and Jin, 2000). The Youjiang Basin was situated at the south of Yangtze block and deposited by extensive deep-water sediments. Shallow-water carbonates developed on some isolated platforms in the Youjiang Basin (Fig. 1B), e.g., Heshan, Leye, and Pingguo (Huang, 1985; Luo and Hou, 1990; Feng et al., 1996; Li et al., 2013). In South China, the LMAs are generally distributed in the Qixia (also termed “Chihsia”) and Maokou formations (Fig. 2; Luo and He, 2010; Liu et al., 2012, 2014; Xue et al., 2015). The Qixia Formation is primarilycomposed of mixed siliciclastic-carbonate successions (close to land), organic matter-rich limestones, limestones bearing cherty nodules (containing LMAs), thick-bedded limestones (containing LMAs), and thin-bedded cherts. The overlying Maokou Formation, from bottom to top, consists of massive LMAs, medium-to thick-bedded limestones episodically interbedded with banded cherts, and alternating shale and silicate deposits (Wang et al., 1994; Feng et al., 1996; Wang and Jin, 2000). Detailed biostratigraphic studies indicate that the LMAs in the Qixia and lower Maokou formations correspond to the Kungurian and Roadian stages, respectively (Fig. 2). A global decrease in sea-level occurred during the Artinskian (Ross and Ross, 1985; Koch and Frank, 2011), which led to the extensive emergence of the Yangtze Platform (Wang and Jin, 2000; Yan et al., 2015; Wu et al., 2016). The following transgression, which occurred between the Artinskian and the Kungurian, facilitated a gradual increase in carbonate deposition on the Yangtze Craton. Afterwards, the sea-level declined in the late Kungurian, and then, it transitioned to a gradual rise in the early Roadian (Wang and Jin, 2000). Subsequently, a marked regression (roughly corresponding to the middle-upper Capitanian stage) occurred, resulting in widely-recognized subaerial exposures on the Yangtze Platform (Fig. 2; He et al., 2005; Yan et al., 2008; Xiao et al., 2016; Haq and Schutter, 2008). Isolated platforms in the Youjiang Basin were also affected by the regression during the late Capitanian (e.g., Bama and Fengshan) (Wang and Shi, 2008; Liu et al., 2014), but they exhibit successive depositional records in relatively deep-water regions (e.g., the Laibin section) (Sha et al., 1990). In this study, we investigated six outcrops and two well sections from different depositional environments (Fig. 1B), including the 766

Marine and Petroleum Geology 111 (2020) 765–785

C. Su, et al.

Fig. 1. (A) Global paleogeography of the Middle Permian. Base map adapted from Scotese (2014). (B) Middle Permian paleogeography of South China (modified from Wang and Jin (2000)) showing study sites and the overall distributions of the limestone−marl alternations (LMAs). Section names: GC2 - Guangcan 2, HY Hongyuan, JK - Jiangkou, L1 - Lai 1, LB - Laibin, LHX - Laohuangxuan, LSX - Lengshuixi, and SS - Shangsi.

platform interior (wells Guangcan 2 and Lai 1; Jiangkou (29° 15′ 29″ N, 107° 51′ 16″ E); Laohuangxuan (29° 28′ 46″ N, 107° 29′ 38″ E); Lengshuixi (29° 53′ 40″ N, 108° 16′ 51″ E)), platform margin (Hongyuan; 31° 32′ 49″ N, 108° 28′ 54″ E), and the shallow-water shelf (Shangsi; 32° 19′ 18″ N, 105° 27′ 22″ E) in the Yangtze Platform, as well as the isolated Heshan platform in the Youjiang Basin (Laibin; 23° 42′ 35″ N, 109° 13′ 47″ E). The LMAs used in this sudy were sampled from both the Qixia and Maokou formations, except for the Laibin section where rocks were only collected from the Qixia Formation (intervals A and B) (Fig. 2). Terrigenous siliciclastics are abundant in the lower Qixia Formation (Interval B) of the Laibin section, but their abundance significantly decreases upwards and reaches a relatively low level in the upper Qixia and lower Maokou formations.

thin-sections and milled powders for further petrological, mineralogical, and geochemical analyses. For each selected LMA, coupled samples were chosen from the limestone and the neighboring marl units. Several transitional layers between the limestone and marl beds were also investigated to assess the potential effect of diagenetic alteration. Petrographic analyses were conducted on 70 polished thinsections using polarized (Leica DM4P) and cathodoluminescence (CL8200 MK5; high voltage of 10 kV, maximum beam of 2 mA, and vacuum of 0.003 mBar) microscopes and on 10 rock cubes using a scanning electron microscopy (SEM; Quanta 650 FEG) to examine the characteristics of their microstructures, including fossils, minerals, and cements. All of these analyses were achieved at the School of Geoscience and Technology, Southwest Petroleum University in Chengdu, China.

3. Material and methods

3.2. X-ray diffraction (XRD) analysis of the mineralogical compositions

3.1. Petrographic analysis

To determine the mineralogical composition of the LMA successions, 35 LMA coupled samples (70 in total) were ground into powders

Bulk rock samples collected from all eight sections were made into 767

Fig. 2. Stratigraphic columns showing the distributions of the LMAs at eight sections. (A) Age and stratigraphic intervals of the Permian strata from the International Commission on Stratigraphy, 2017(http://www. stratigraphy.org/). Conodont zonations modified from Shen et al. (2019) (Sw. = Sweetognathus, J. = Jinogondolella, and C. = Clarkina). (B) Lithological log of the Shangsi section. Thickness (thick.) data from Yan et al. (2008) and Ma et al. (2008), and conodont data from Sun et al. (2008) and Fang et al. (2012). (C)–(H) Logs of the Hongyuan, Guangcan 2, Lai 1, Lengshuixi, Laohuangxuan and Jiangkou sections. Note that natural gamma-ray (GR) curves were used for stratigraphic correlation. (I) Log of the Laibin section. Thickness data from Sha et al. (1990), and conodont data from Shen et al. (2007, 2019).

C. Su, et al.

Marine and Petroleum Geology 111 (2020) 765–785

768

Marine and Petroleum Geology 111 (2020) 765–785

C. Su, et al.

Fig. 3. Lithological characteristics of the LMAs observed in the field. (A) Irregularly-bedded LMA, interval A in the Laibin section. The ruler is 20 cm in length. (B) Well-bedded mixed carbonate-siliciclastic successions with sharp boundaries between the limestone and marl beds, interval B in the Laibin section. The pen is 14 cm in length. (C) Cut surface of the yellow box in (A), showing the transitional zone between the marl and limestone beds. (D) Nodular to eyeball-like (elliptical) limestone beds in the LMAs at Shangsi section. Yellow arrows indicate two nearly fully dissolved limestone beds. (E) A large well-formed alatoconchid bivalve in a limestone bed in an LMA. Lower Maokou Formation in the Shangsi section. (F) Bioclasts exhibiting their primary shapes in a limestone bed (yellow arrows). The length of the pen is 14 cm. (G) Fossils in a marl bed exhibiting strongly compacted and oriented features (yellow arrows) compared to the adjacent limestone (F). Both (F) and (G) are from the Laohuangxuan section. Note L - limestone, M - marlstone, T - transitional zone, Se - sepiolite, and Ga - gastropod. (For interpretation of the references to color in this figure legend, the reader is referred to the Web version of this article.)

769

Marine and Petroleum Geology 111 (2020) 765–785

C. Su, et al.

(caption on next page) 770

Marine and Petroleum Geology 111 (2020) 765–785

C. Su, et al.

Fig. 4. Photomicrographs of the LMAs in the Permian strata of South China. (A) Skeletal fragments preserved in marl and limestone beds. Interval A at Laibin section. (B) Enlargement of the upper box in (A) showing the abundant well-formed fossils in the limestone bed. (C) Paired cathodoluminescence image of (B). (D) Enlargement of the lower box in (A). Numerous skeletal fragments oriented roughly parallel to the bedding surface. (E) Paired cathodoluminescence image of (D). Dark red cathodoluminescence of the matrix. (F) Limestone bed (mudstone) in an LMA in interval B at Laibin section containing rare pyrite grains. (G) Marl bed with high volumes of bioclasts (see (H) and (I)). Interval B in the Laibin section. (H) Enlargement of the yellow box in (G). Fossils including bivalve and bryozoan fragments and foraminifer in a marl bed exhibiting negligible compaction. (I) Paired cathodoluminescence image of (H) showing large quantities of bioclasts preserved in a marl bed. (J) A polished LMA slab from the Shangsi section clearly showing the marl, limestone, and transitional beds. Note the large number of sepiolite minerals currently preserved in each bed. (K) Enlargement of the yellow box in (J). Strongly oriented small fossil fragments and calcite veins parallel to the bedding surface and abundant brown sepiolite (Se) in the matrix. (L) Paired cathodoluminescence image of (K) showing numerous bright, euhedral dolomite crystals. (M) and (N) Preserved characteristics of Dasycladalean algae in the limestone and marl beds. Note that strong compaction occurred in marl the bed (N). Laohuangxuan section. (O) SEM image showing the authigenic sepiolite with fibrous fabrics. Shangsi section. (P) Transformation from sepiolite to talc preserved in the marl beds of the LMAs. Hongyuan Section. (Q) Illite preserved within intergranular voids showing some alteration and deformation characteristics. Interval B in the Laibin section. (For interpretation of the references to color in this figure legend, the reader is referred to the Web version of this article.)

with particle sizes of 200 mesh and finer for XRD analysis (PANalytical X'Pert MPD PRO) at the State Key Laboratory of Oil and Gas Reservoir Geology and Exploration, Southwest Petroleum University. A Rigaku D/ max-rA diffractometer equipped with a Cu Kα radiation source was used to record the mineral XRD patterns. The mineralogical characteristics of the samples were compared with the Joint Committee Powder Diffraction Standards database, and the results were normalized by mass. For samples with low clay mineral contents, the samples were separated by suspension extraction, and then, they were processed using three methods before analysis, i.e., naturally dried, ethylene glycol, and high-temperature (550 °C) methods (Eslinger and Pevear, 1988).

components (relying on cathodoluminescence examination), and (2) to avoid the possible clay REE fraction using single-step acid digestion (Tostevin et al., 2016; Li et al., 2019). In situ LA analysis was performed using a Coherent GeoLasPro system (193 nm COMPexPro 102 ArF excimer laser) connected to an Agilent 7700e ICP-MS instrument at the Sample-Solution Analytical Laboratory in Wuhan, China. The spot size and laser frequency were set to 44 μm and 5 Hz, respectively (Li et al., 2017, 2019). The elemental compositions of the minerals were calibrated using the internal independent standard strategy (Liu et al., 2008; Chen et al., 2011). Other procedures follow those described by Li et al. (2019). The measured REE + Y isotopes include 89Y, 139La, 140Ce, 141 Pr, 146Nd, 147Sm, 153Eu, 157Gd, 159Tb, 163Dy, 165Ho, 166Er, 169Tm, 172 Yb, and 175Lu. Their relative deviations were less than 9% by monitoring the carbonate standard MACS-3 (n = 36) used in each sequence (Table S1). Several other isotopes (27Al, 49Ti, 57Fe, 91Zr, and 232Th) were also measured with relative deviations of < 10% (Table S1).

3.3. Element geochemistry 3.3.1. X-ray fluorescence (XRF) analysis of major elements Twenty groups of LMA successions (40 samples) of coupled limestone and marl beds from the Hongyuan, Jiangkou, Laohuangxuan, and Lengshuixi sections (Eastern Sichuan Basin (ESB)) were ground into powders (~200 mesh) for major element analysis, including Al2O3, TiO2, and Fe2O3. Approximately 1.0 g of dried powder from each sample was placed in a the ceramic crucible and heated in a muffle furnace (Hooper, 1969). Then, the melted material was quenched and made into a flat disc, which was analyzed on the XRF (ZSX Primus Ⅱ) at the Sample-Solution Analytical Laboratory in Wuhan, China. The reliability of the results was monitored using two standards (GBW07108 and GBW07114) and four repeated samples, resulting in a measurement uncertainty of less than 10%.

4. Results 4.1. Petrographic and sedimentological characteristics Morphologically, the limestone beds (commonly 10–30 cm in thickness) show normal to irregular bedding and nodular structures within the LMAs, while the marl beds (< 15 cm in thickness) are generally parallel to or warp around the neighboring limestone units (Fig. 3). The transitional layer between the coupled limestone and marl beds of the shallow-water LMA successions, which contain minimal terrestrial material, is mostly unremarkable (Fig. 3A, C, and F). Some of the boundaries exhibit mm-scale spatial variations in color, fossil preservation, and mineralogy (Fig. 4A and J). However, a clear-cut boundary between the limestone and the marl beds has commonly been identified in interval B of the lower Qixia Formation at Laibin section (Figs. 2I and 3B). The limestone and marl beds within the LMA succession exhibit recognizable petrographic differences. In the siliciclastic-rich LMAs from interval B of the Laibin section, the marl beds are characterized by bioclast-dominant (~55%) packstone textures, unclear orientations, and relatively high organic matter contents (Fig. 4G–I). In comparison, the limestone beds consist of mudstone, rare fossils (< 5%), and a few pyrite grains (< 3%) (Fig. 4F). In the siliciclastic-poor LMAs, examination of the thin-sections reveals that the fossil contents of the limestone beds are generally lower than those of the adjacent marl beds (Fig. 4A and J). Fragments of calcareous algae, brachiopod, ostracod, echinoderm, bryozoan, and foraminifer are plentiful in both the limestone and marl beds (Fig. 3E–G and Fig. 4A–G), but some large fossils (e.g., a giant alatoconchid bivalve with a ~20 -cm-long shell) and primary aragonite skeletons (e.g., gastropods) are well preserved in the limestone beds (Fig. 3E and F). Fossils in the marl beds show markedly increased deformation structures and preferred orientations (Fig. 3G and 4N) compared to the limestone beds (Figs. 3F and 4M). Furthermore, the matrix of the limestone beds is generally composed of microspar calcite crystals (Fig. 4B, C, and M), while the matrix of the marl

3.3.2. Solution-based element analysis Ten sets of bulk rock samples from the Laibin and Shangsi sections were measured using inductively coupled plasma-optical emission spectrometry (ICP-OES). Then, the ICP-OES results were compared to the XRF results. The detailed procedures are described by Olesik (1991). In addition, 10 sets of bulk rock samples from the Laibin and Shangsi sections and other 20 sets of bulk samples (same as samples for XRF) from Hongyuan, Jiangkou, Laohuangxuan, and Lengshuixi were digested in an HF + HNO3 solution and analyzed using an ICP-MS (Agilent 7700e) to determine their trace and rare earth element concentrations. To ensure the reliability of the data, we measured 10 samples twice. Ten blanks and 20 standards (AGV-2 × 5, BHVO-2 × 5, BCR-2 × 5, and RGM-2 × 5) were also analyzed in sequence to control the accuracy of the measurements, producing relative errors of less than 5%. The major-element ICP-OES analyses were performed at the Chengdu University of Technology; the trace and rare earth element ICP-MS analyses were performed at the Sample-Solution Analytical Laboratory in Wuhan, China. 3.3.3. In situ LA-ICP-MS analysis of the coupled LMA successions The LA-ICP-MS analysis was conducted (1) to obtain precise results for the limestone and marl components within the LMA by excluding interferences from veins, fossils, and the strongly diagenetic 771

Marine and Petroleum Geology 111 (2020) 765–785

C. Su, et al.

Table 1 The mineralogical compositions of the LMA samples within the sampling interval in the field sections. Section names: HY - Hongyuan, JK - Jiangkou, LBA - Interval A of Laibin, LBB - Interval B of Laibin, LHX - Laohuangxuan, LSX - Lengshuixi), and SS - Shangsi. Sample NO.

Sample location (m)

Lithology

Percentage content × 10−2 Carbonates

SS-1-1 SS-1-2 SS-2-1 SS-2-2 SS-3-1 SS-3-2 SS-4-1 SS-4-2 SS-5 SS-6 SS-7-1 SS-7-2 SS-8-1 SS-8-2 SS-9-1 SS-9-2 SS-10-1 SS-10-2 SS-11-1 SS-11-2 HY-1-1 HY-1-2 HY-2-1 HY-2-2 HY-3-1 HY-3-2 HY-4-1 HY-4-2 LSX-1-1 LSX-1-2 LSX-2-1 LSX-2-2 JK-1-1 JK-1-2 JK-2-1 JK-2-2 JK-3-1 JK-3-2 JK-4-1 JK-4-2 JK-5-1 JK-5-2 JK-6-1 JK-6-2 JK-7-1 JK-7-2 JK-8-1 JK-8-2 LHX-1-1 LHX-1-2 LHX-2-1 LHX-2-2 LHX-3-1 LHX-3-2 LHX-4-1 LHX-4-2 LHX-5-1 LHX-5-2 LHX-6-1 LHX-6-2 LBA-1-1 LBA-1-2 LBA-5-1 LBA-5-2 LBA-6-1 LBA-6-2 LBB-7-1 LBB-7-2 LBB-8-1 LBB-8-2

147.20 147.24 147.70 147.73 148.20 148.23 148.70 148.72 149.20 149.70 98.10 98.12 103.20 103.23 109.60 109.62 115.75 115.77 122.50 122.52 99.50 99.53 118.20 118.22 135.00 135.02 147.33 147.35 126.33 126.35 180.70 180.72 102.21 102.24 110.56 110.58 117.33 117.35 127.02 127.04 133.41 133.44 141.10 141.12 148.90 148.92 157.30 157.32 95.10 95.12 113.52 113.55 135.40 135.43 155.62 155.64 177.12 177.14 192.50 192.52 343.11 343.13 342.00 342.02 258.49 258.52 22.80 22.82 22.62 22.63

Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Marl Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone Marl Limestone

Quartz

Calcite

Dolomite

26 82 25 95 29 88 30 90 23 18 92 95 99 97 81 98 94 98 66 90 50 75 97 99 82 99 89 98 88 99 96 97 92 98 63 98 76 93 84 98 75 98 70 97 88 99 85 98 96 88 90 99 98 84 83 92 89 98 96 99 25 63 79 94 86 100 43 65 31 95

2 1 5 5 1 3 13 6 5 3 2 18 4 40 23

Sepiolite 6 2 5 1 8 2 6 1 9 9 2

3 1 61 30 9 3

Montmorillonite

Illite

Clinochlore

9 7 3 16 2 10

11 6 2 16 6 3 2 1

2 1 2 1 3 1 1 1 7 1 13 1 8 1 11 1 4 2 4 6 1 2 3 3

3

57 16 62 4 55 10 43 8 63 50

Talcum

1

3

2

Clay minerals

1 6 4 4 6

14 7 12 3 14 29 24 46 5

772

8

7 2 1

15 1 9 1

10

1

10 4 7 1 4 10 6 1 6 1 3 1 4 1 4 1

20 1 7 6 1

7 1

10 1 16 1 2 3 9 4 1 7 9 1 3 1

2

1

5 7

8 5 9

15 6 7

Marine and Petroleum Geology 111 (2020) 765–785

C. Su, et al.

Fig. 5. Ternary plots showing the mineral compositions of the LMAs. Carbonates include calcite and dolomite. Sepiolite and talc are the dominant clay minerals. The three samples marked by black arrows are from interval B in the Laibin section and contain > 20% detrital quartz. Marl samples with a clay mineral content surpassing 50% are circled in red. These samples are from the Maokou Formation at Shangsi section and are dominated by sepiolite. ESB Eastern Sichuan Basin, including the Hongyuan, Jiangkou, Laohuangxuan, and Lengshuixi sections. (For interpretation of the references to color in this figure legend, the reader is referred to the Web version of this article.)

beds is composed of micrite, clay minerals, and minor to rare siliciclastics (Fig. 4D, K, and N).

interval B of the Laibin section have much higher Al2O3 (mean ± SD = 5.15 ± 2.48 wt%; n = 4), Th (5.53 ± 2.51 ppm), Zr (100.1 ± 49.2 ppm), and ΣREE (79.9 ± 21.8 ppm) contents than the other sites (Al2O3 = 0.36 ± 0.35 wt%, Th = 0.52 ± 0.66 ppm, Zr = 6.5 ± 6.2 ppm, ΣREE = 7.2 ± 5.7 ppm; n = 56). In addition, in all eight sections, the marl beds had significantly higher Al, Th, Zr, and ΣREE contents than their neighboring limestone beds (Table 2). The analytical results of the LA method are similar to the solution-based bulk rock analyses (Tables 2 and S2), but the values acquired using the LA method are generally slightly lower than the solution analyses (Table 3).

4.2. Mineralogy The mineralogical compositions were determined using XRD analysis (for original data, see Table 1) and the results are plotted on a ternary diagram (Fig. 5). Fig. 5 shows that the LMA are primarily composed of calcite, dolomite, quartz, and clays. Calcite is the dominant mineral in both the limestone and marl beds of the LMAs (Table 1). The dolomite is composed of fine-grained, subhedral to euhedral crystals (Fig. 4L). In the analyzed LMA successions at all of the study sites, the marl beds (n = 36) generally have higher volumes of quartz (mean value of 8%) and clay minerals (17%) than adjacent limestone beds (2% and 2%, respectively; n = 34). The clay minerals are principally composed of sepiolite and talc (Fig. 4O and P) except for the samples from the interval B of Laibin section, which consists of montmorillonite, illite, and clinochlore (Table 1, Fig. 4Q). Microscopic examination shows that sepiolite and talc have partially or completely replaced the bioclasts/matrix (Fig. 4J, K, O, and P). The quartz content is generally low in all of the LMA sections (< 16%) except for the interval B of Laibin section (26%; n = 4). Morphologically, the quartz grains (< 0.2 mm in size) are subangular to subrounded in interval B of Laibin section (Fig. 4G and H). However, at all of the other sites, the quartz is generally present as nodular chert or as the replacement product of fossils (e.g., Fig. 4M).

5. Discussion 5.1. Primary depositional LMA successions: Interval B of the Qixia Formation at Laibin section The lithology, mineralogy, and element geochemical characteristics indicate that the formation of LMAs in interval B of the Qixia Formation was principally controlled by the original depositional process. First, the LMAs do not show signs of differential compaction caused by diagenesis (e.g., the dissolution traits in Fig. 4D, K, and N), although they do exhibit physical compaction features (Fig. 4F–I). Skeletal grains are rare in the limestone beds (generally mudstones) but are abundant and well-preserved in the marl beds (Fig. 4F–I). Second, the XRD and petrographic results show much higher proportions of the non-carbonate components in the marls than in the adjacent limestones (Fig. 5; Table 1). The components of these siliciclastic and clay minerals include quartz, montmorillonite, illite, and chlorite (Fig. 4Q). The assemblage of non-carbonate minerals and their morphological features, i.e., signs of transportation and abrasion, indicate that the minerals are the result of terrigenous input rather than authigenic minerals formed during diagenesis (see Table 1 and Section 4.2) (Chamley, 1998). Third, several elements (e.g., Al, Zr, and REEs) are sensitive to the admixture of exogenetic siliciclastic components in the carbonates (Kamber et al., 2004; Li et al., 2019), and thus, they can be used to evaluate of detrital supply during the generation of the LMAs. Both bulk rock and in situ

4.3. Element compositions of the LMAs The element concentration data for the limestone and marl components of the bulk LMAs are reported in Table 2. The mean values of the different components from of each spot acquired using in situ LAICP-MS are also listed in Table 2, and the original data for the spots are listed in Table S2. Al and Ti were normalized to Al2O3 and TiO2, respectively, for comparison. The LMAs were separated into two groups based on their major and trace element compositions. Both the limestone and marl beds in 773

Sample location (m)

99.53 118.22 118.20 135.02 135.00 147.35 147.33 102.24 102.21 110.58 110.56 117.35 117.33 127.04 127.02 133.44 133.41 141.12 141.10 148.92 148.90 157.32 157.30 343.13 343.11 342.02 342.00 258.52 258.49 22.82 22.80 22.63 22.62 95.12 95.10 113.55 113.52 135.43 135.40 155.64 155.62 177.14 177.12 192.52 192.50 126.35 126.33 180.72 180.70 147.24 147.20 147.73 147.70

Sample No.

B-HY-1L B-HY-2L B-HY-2M B-HY-3L B-HY-3M B-HY-4L B-HY-4M B-JK-1L B-JK-1M B-JK-2L B-JK-2M B-JK-3L B-JK-3M B-JK-4L B-JK-4M B-JK-5L B-JK-5M B-JK-6L B-JK-6M B-JK-7L B-JK-7M B-JK-8L B-JK-8M B-LBA-1L B-LBA-1M B-LBA-5L B-LBA-5M B-LBA-6L B-LBA-6M B-LBB-7L B-LBB-7M B-LBB-8L B-LBB-8M B-LHX-1L B-LHX-1M B-LHX-2L B-LHX-2M B-LHX-3L B-LHX-3M B-LHX-4L B-LHX-4M B-LHX-5L B-LHX-5M B-LHX-6L B-LHX-6M B-LXS-1L B-LXS-1M B-LXS-2L B-LXS-2M B-SS-1L B-SS-1M B-SS-2L B-SS-2M

Table 2

0.205 0.060 0.365 0.120 0.419 0.149 0.674 0.093 0.796 0.072 0.446 0.111 0.505 0.207 0.390 0.194 0.395 0.142 0.621 0.087 0.432 0.112 0.392 0.133 0.304 0.095 0.254 0.424 1.783 4.746 7.943 1.342 6.587 0.163 0.914 0.078 0.324 0.172 0.109 0.278 0.333 0.293 0.313 0.149 1.688 0.060 1.091 0.071 0.246 0.208 0.633 0.101 0.640

Al2O3 (wt %) 0.0160 0.0110 0.0190 0.0100 0.0220 0.0120 0.0420 0.0100 0.0220 0.0080 0.0290 0.0090 0.0410 0.0160 0.0300 0.0050 0.0310 0.0150 0.0270 0.0130 0.0230 0.0090 0.0240 0.0120 0.0160 0.0090 0.0140 0.0230 0.0750 0.1670 0.2950 0.0500 0.2380 0.0120 0.0510 0.0080 0.0200 0.0160 0.0100 0.0240 0.0220 0.0130 0.0170 0.0110 0.0770 0.0070 0.0510 0.0080 0.0150 0.0130 0.0420 0.0090 0.0440

TiO2 (wt%)

476 133 833 119 966 168 1904 203 1001 42 1064 112 1813 322 1127 98 1085 455 1120 168 770 308 952 470 2218 424 750 1512 4349 11669 19698 4365 18578 238 2142 161 546 406 119 980 854 259 861 679 6272 56 3031 / 364 480 1831 264 1863

Fe (ppm)

0.15 0.06 0.24 0.07 0.36 0.12 0.61 0.07 4.26 0.05 0.33 0.07 0.48 0.14 0.30 0.15 0.39 0.18 0.41 0.06 0.25 0.11 0.34 0.58 0.77 0.54 0.70 0.97 2.11 5.10 8.36 1.67 7.01 0.09 0.99 0.07 0.20 0.14 0.06 0.24 0.34 0.26 0.67 0.31 1.51 0.02 1.58 0.02 0.11 0.60 0.92 0.67 1.08

Th (ppm)

2.37 1.00 4.00 1.05 4.93 1.62 8.21 1.25 12.21 0.99 4.56 1.18 7.21 2.40 4.56 2.28 4.99 1.99 4.27 1.28 3.38 1.56 5.23 7.74 13.22 7.76 9.82 12.31 33.01 93.41 161.90 26.16 119.20 1.55 12.93 1.71 3.07 2.84 1.10 4.07 4.54 2.04 4.94 1.56 20.43 0.61 13.27 0.41 2.14 8.47 17.90 7.22 18.53

Zr (ppm)

1.314 0.506 1.415 1.834 2.914 2.448 4.033 1.096 4.816 0.508 3.912 0.747 1.168 0.919 1.457 1.645 1.179 2.126 1.939 1.475 2.556 3.122 4.433 3.946 6.066 4.605 4.873 7.030 11.290 14.580 19.250 11.120 15.770 1.479 1.609 0.548 0.924 1.151 0.551 2.502 2.983 1.769 0.929 1.721 2.825 0.779 3.426 0.150 0.838 2.817 3.449 2.178 3.929

Y (ppm)

0.655 0.323 0.804 0.903 1.205 1.407 2.151 0.718 3.303 0.382 6.796 0.436 1.132 0.938 0.691 0.959 0.890 2.353 1.832 0.662 1.816 1.023 1.284 2.846 3.874 1.230 2.005 3.400 7.844 19.800 28.260 12.210 22.680 0.851 0.792 0.325 0.751 0.553 0.159 0.940 0.851 1.917 0.400 1.394 2.561 0.212 3.301 0.082 0.450 1.922 1.723 0.910 1.703

La (ppm)

1.201 0.596 1.716 1.351 2.208 1.926 4.223 1.231 7.308 0.519 1.850 0.793 2.458 1.265 1.382 1.684 1.951 3.499 2.264 1.018 2.096 1.474 2.076 3.187 4.221 2.552 3.417 5.372 11.060 26.060 36.110 15.690 34.420 1.246 1.774 0.601 1.495 1.223 0.333 1.696 1.814 3.873 0.970 2.975 5.505 0.265 8.086 0.102 0.935 3.980 4.037 2.564 4.110

0.147 0.075 0.188 0.160 0.281 0.237 0.554 0.139 0.830 0.074 1.292 0.080 0.264 0.173 0.161 0.179 0.213 0.430 0.304 0.121 0.340 0.187 0.259 0.525 0.713 0.323 0.497 0.641 1.493 3.560 4.579 2.118 4.747 0.164 0.184 0.072 0.163 0.136 0.041 0.201 0.205 0.464 0.114 0.363 0.625 0.038 0.779 0.015 0.116 0.460 0.471 0.298 0.476

Pr (ppm)

(continued on next page)

Ce (ppm)

C. Su, et al.

Marine and Petroleum Geology 111 (2020) 765–785

774

775

Nd (ppm)

0.540 0.262 0.745 0.665

B-HY-1L B-HY-2L B-HY-2M B-HY-3L

148.23 148.20 148.72 148.70 149.20 149.70 140.00 140.01 140.01 135.01 135.00 135.00 75.10 75.11 75.11 343.13 343.12 342.71 342.71 342.71 342.45 342.45 342.45 342.22 342.22 342.22 342.00 342.00 342.00 258.51 258.50 22.82 22.80 22.63 22.62 147.73 147.70 148.23 148.20 148.72 148.70 148.71 149.20 149.70

B-SS-3L B-SS-3M B-SS-4L B-SS-4M B-SS-5M B-SS-6M I-GC2-1L I-GC2-1M I-GC2-1T I-HY-3L I-HY-3M I-HY-3T I-L1-1L I-L1-1M I-L1-1T I-LBA-1L I-LBA-1M I-LBA-2L I-LBA-2M I-LBA-2T I-LBA-3L I-LBA-3M I-LBA-3T I-LBA-4L I-LBA-4M I-LBA-4T I-LBA-5L I-LBA-5M I-LBA-5T I-LBA-6L I-LB-6M I-LBB-7L I-LBB-7M I-LBB-8L I-LBB-8M I-SS-2L I-SS-2M I-SS-3L I-SS-3M I-SS-4L I-SS-4M I-SS-4T I-SS-5M I-SS-6M

Sample No.

Sample location (m)

Sample No.

Table 2 (continued)

0.110 0.052 0.141 0.153

Sm (ppm)

0.136 0.456 0.089 0.347 0.537 0.335 0.014 0.342 0.193 0.055 0.368 0.094 0.141 0.105 0.173 0.022 0.105 0.048 0.254 0.024 0.020 0.172 0.024 0.115 0.282 0.063 0.031 0.098 0.104 0.462 1.407 4.241 4.772 0.911 3.284 0.042 0.398 0.051 0.340 0.065 0.157 0.089 0.405 0.222

Al2O3 (wt %)

0.020 0.009 0.038 0.030

Eu (ppm)

0.0110 0.0250 0.0090 0.0200 0.0320 0.0180 0.0010 0.0140 0.0010 0.0020 0.0330 0.0020 0.0040 0.0040 0.0030 0.0040 0.0100 0.0002 0.0130 0.0002 0.0002 0.0250 0.0004 0.0060 0.0110 0.0030 0.0020 0.0040 0.0050 0.0170 0.0430 0.1390 0.1580 0.0510 0.1040 0.0020 0.0420 0.0020 0.0110 0.0020 0.0080 0.0030 0.0210 0.0080

TiO2 (wt%)

0.129 0.047 0.170 0.148

Gd (ppm)

426 1224 603 1169 1810 872 110 618 68 90 938 126 77 292 113 183 342 172 306 356 122 235 262 138 1233 138 94 176 291 941 1338 5131 5656 1745 6719 131 634 154 391 153 201 404 730 231

Fe (ppm)

0.020 0.007 0.031 0.024

Tb (ppm)

0.57 0.75 0.50 0.69 0.83 0.68 0.04 0.19 0.15 0.07 0.19 0.09 0.24 0.50 0.50 0.06 0.27 0.03 0.28 0.10 0.04 0.14 0.04 0.04 0.29 0.10 0.10 0.23 0.19 0.43 0.76 3.86 5.04 0.77 4.83 0.04 0.34 0.04 0.17 0.04 0.13 0.07 0.26 0.10

Th (ppm)

0.112 0.055 0.160 0.169

Dy (ppm)

7.79 13.13 6.90 11.17 15.13 10.67 0.67 2.82 2.99 0.47 2.81 0.91 0.51 0.61 0.91 0.56 1.89 0.32 2.53 0.33 0.54 2.81 0.38 1.08 5.95 0.26 0.50 1.16 3.15 3.29 12.29 43.73 45.35 9.62 29.94 0.42 6.22 0.63 3.79 0.65 1.64 0.67 4.07 3.45

0.027 0.011 0.043 0.031

Ho (ppm)

Zr (ppm)

0.080 0.032 0.107 0.103

Er (ppm)

2.428 3.363 2.053 3.284 3.725 3.331 0.898 1.171 1.007 2.186 2.456 2.538 0.607 0.936 0.860 2.404 4.589 2.638 3.820 4.649 1.802 2.557 1.815 1.561 2.929 2.260 4.537 4.198 4.203 5.267 5.101 10.500 13.790 9.471 12.520 0.706 1.873 0.896 1.543 0.910 1.391 1.072 1.707 0.956

Y (ppm)

0.012 0.004 0.019 0.016

Tm (ppm)

1.136 1.100 0.595 1.246 1.465 1.150 0.440 0.494 0.580 1.426 1.249 1.653 1.378 1.921 1.594 2.127 2.747 1.026 1.932 2.237 0.320 0.370 0.342 0.291 0.486 0.338 1.266 1.290 1.228 2.237 3.790 12.110 16.630 9.381 13.940 0.462 1.233 0.349 0.544 0.515 0.692 0.702 0.626 0.365

La (ppm)

0.081 0.027 0.086 0.081

0.012 0.005 0.018 0.011

Lu (ppm)

0.331 0.360 0.249 0.378 0.441 0.361 0.071 0.116 0.136 0.256 0.232 0.294 0.300 0.412 0.348 0.327 0.450 0.165 0.391 0.364 0.053 0.076 0.063 0.054 0.151 0.086 0.186 0.211 0.191 0.371 0.856 2.354 3.984 1.635 3.962 0.092 0.299 0.076 0.146 0.097 0.149 0.134 0.142 0.085

Pr (ppm)

(continued on next page)

Yb (ppm)

2.866 3.109 2.173 3.220 3.702 3.175 0.623 0.924 1.101 1.906 1.634 2.251 2.570 3.679 3.162 1.335 2.240 1.052 2.431 2.130 0.435 0.562 0.515 0.478 1.000 0.640 1.163 1.348 1.239 3.193 6.387 19.410 28.000 11.770 27.690 0.718 2.504 0.559 1.109 0.881 1.050 1.183 1.291 0.771

Ce (ppm)

C. Su, et al.

Marine and Petroleum Geology 111 (2020) 765–785

Nd (ppm)

1.129 0.947 2.139 0.463 3.081 0.293 4.731 0.303 0.976 0.633 0.601 0.655 0.728 1.706 1.166 0.554 1.282 0.692 1.040 2.878 3.609 1.916 2.623 3.219 6.724 14.740 18.200 9.096 19.470 0.641 0.706 0.277 0.609 0.481 0.135 0.784 0.876 1.659 0.474 1.456 2.364 0.156 2.967 0.049 0.393 2.340 2.461 1.798 2.489 1.922 2.002 1.613 2.099 2.320

Sample No.

B-HY-3M B-HY-4L B-HY-4M B-JK-1L B-JK-1M B-JK-2L B-JK-2M B-JK-3L B-JK-3M B-JK-4L B-JK-4M B-JK-5L B-JK-5M B-JK-6L B-JK-6M B-JK-7L B-JK-7M B-JK-8L B-JK-8M B-LBA-1L B-LBA-1M B-LBA-5L B-LBA-5M B-LBA-6L B-LBA-6M B-LBB-7L B-LBB-7M B-LBB-8L B-LBB-8M B-LHX-1L B-LHX-1M B-LHX-2L B-LHX-2M B-LHX-3L B-LHX-3M B-LHX-4L B-LHX-4M B-LHX-5L B-LHX-5M B-LHX-6L B-LHX-6M B-LXS-1L B-LXS-1M B-LXS-2L B-LXS-2M B-SS-1L B-SS-1M B-SS-2L B-SS-2M B-SS-3L B-SS-3M B-SS-4L B-SS-4M B-SS-5M

Table 2 (continued)

0.219 0.241 0.391 0.131 0.694 0.062 0.818 0.076 0.193 0.125 0.120 0.198 0.140 0.370 0.229 0.124 0.260 0.181 0.225 0.556 0.758 0.361 0.509 0.628 1.375 2.833 3.503 1.647 3.639 0.167 0.174 0.051 0.145 0.105 0.048 0.212 0.213 0.378 0.119 0.288 0.477 0.027 0.583 0.024 0.080 0.409 0.401 0.298 0.427 0.324 0.319 0.255 0.382 0.409

Sm (ppm) 0.042 0.058 0.074 0.021 0.031 0.015 0.156 0.019 0.033 0.028 0.026 0.026 0.025 0.078 0.041 0.024 0.058 0.045 0.042 0.127 0.168 0.088 0.109 0.149 0.303 0.554 0.654 0.401 0.599 0.030 0.072 0.008 0.028 0.023 0.002 0.043 0.043 0.031 0.015 0.054 0.090 0.010 0.081 0.020 0.019 0.086 0.089 0.069 0.095 0.066 0.072 0.054 0.073 0.079

Eu (ppm) 0.204 0.277 0.320 0.096 0.697 0.048 0.674 0.075 0.137 0.134 0.101 0.174 0.101 0.310 0.229 0.127 0.263 0.264 0.326 0.413 0.630 0.298 0.411 0.561 1.317 2.498 3.087 1.401 2.971 0.169 0.145 0.053 0.071 0.076 0.029 0.223 0.222 0.281 0.117 0.278 0.418 0.047 0.443 0.021 0.074 0.207 0.223 0.107 0.240 0.107 0.141 0.065 0.147 0.178

Gd (ppm) 0.038 0.041 0.066 0.016 0.125 0.008 0.097 0.012 0.026 0.018 0.020 0.028 0.020 0.056 0.032 0.019 0.034 0.047 0.057 0.066 0.109 0.060 0.073 0.094 0.215 0.385 0.491 0.219 0.456 0.022 0.030 0.007 0.015 0.015 0.006 0.035 0.039 0.045 0.023 0.034 0.072 0.009 0.079 0.002 0.013 0.042 0.048 0.027 0.056 0.028 0.036 0.021 0.036 0.043

Tb (ppm) 0.247 0.253 0.385 0.097 0.746 0.041 0.488 0.060 0.139 0.113 0.132 0.160 0.094 0.262 0.173 0.125 0.226 0.277 0.398 0.424 0.643 0.412 0.507 0.667 1.381 2.262 2.948 1.274 2.547 0.138 0.197 0.042 0.077 0.100 0.031 0.231 0.252 0.262 0.124 0.247 0.399 0.061 0.437 0.012 0.082 0.275 0.344 0.187 0.370 0.204 0.243 0.155 0.283 0.332

Dy (ppm) 0.060 0.059 0.092 0.022 0.155 0.011 0.091 0.016 0.034 0.025 0.031 0.036 0.026 0.057 0.042 0.030 0.053 0.066 0.093 0.103 0.158 0.109 0.130 0.157 0.299 0.457 0.614 0.283 0.510 0.031 0.042 0.012 0.020 0.025 0.009 0.054 0.059 0.054 0.029 0.044 0.084 0.011 0.102 0.003 0.018 0.075 0.093 0.058 0.106 0.065 0.080 0.055 0.082 0.094

Ho (ppm) 0.178 0.151 0.255 0.055 0.426 0.038 0.249 0.041 0.092 0.062 0.098 0.097 0.088 0.142 0.133 0.080 0.146 0.175 0.297 0.210 0.351 0.249 0.311 0.391 0.806 1.275 1.712 0.739 1.388 0.075 0.150 0.031 0.060 0.066 0.024 0.151 0.176 0.150 0.093 0.146 0.248 0.040 0.287 0.008 0.057 0.154 0.232 0.107 0.260 0.124 0.199 0.093 0.182 0.236

Er (ppm) 0.028 0.023 0.037 0.007 0.064 0.004 0.032 0.009 0.017 0.011 0.016 0.016 0.012 0.022 0.017 0.011 0.019 0.024 0.036 0.032 0.054 0.040 0.046 0.056 0.114 0.188 0.252 0.103 0.196 0.010 0.025 0.005 0.008 0.013 0.004 0.028 0.024 0.023 0.012 0.023 0.041 0.005 0.045 0.001 0.009 0.028 0.038 0.021 0.042 0.023 0.034 0.018 0.035 0.039

Tm (ppm) 0.161 0.109 0.234 0.046 0.419 0.025 0.192 0.039 0.093 0.055 0.089 0.081 0.073 0.130 0.121 0.063 0.110 0.135 0.209 0.300 0.409 0.339 0.390 0.433 0.760 1.247 1.725 0.751 1.373 0.060 0.184 0.031 0.059 0.075 0.027 0.147 0.155 0.144 0.098 0.121 0.251 0.039 0.289 0.010 0.059 0.278 0.367 0.235 0.399 0.256 0.327 0.218 0.310 0.361

0.021 0.020 0.039 0.007 0.055 0.005 0.029 0.007 0.018 0.008 0.014 0.013 0.015 0.018 0.014 0.010 0.016 0.020 0.032 0.037 0.058 0.045 0.051 0.061 0.107 0.195 0.271 0.116 0.207 0.009 0.026 0.003 0.010 0.013 0.006 0.023 0.026 0.021 0.016 0.018 0.044 0.006 0.046 0.001 0.010 0.034 0.050 0.027 0.049 0.031 0.043 0.026 0.042 0.050

Lu (ppm)

(continued on next page)

Yb (ppm)

C. Su, et al.

Marine and Petroleum Geology 111 (2020) 765–785

776

Nd (ppm)

2.052 0.313 0.452 0.581 0.990 0.917 1.236 1.064 1.495 1.315 1.468 1.930 0.668 1.600 1.300 0.216 0.334 0.263 0.228 0.537 0.331 0.688 0.840 0.783 1.525 3.442 9.213 16.610 6.660 16.720 0.394 1.192 0.290 0.476 0.367 0.563 0.494 0.534 0.323

Sample No.

B-SS-6M I-GC2-1L I-GC2-1M I-GC2-1T I-HY-3L I-HY-3M I-HY-3T I-L1-1L I-L1-1M I-L1-1T I-LBA-1L I-LBA-1M I-LBA-2L I-LBA-2M I-LBA-2T I-LBA-3L I-LBA-3M I-LBA-3T I-LBA-4L I-LBA-4M I-LBA-4T I-LBA-5L I-LBA-5M I-LBA-5T I-LBA-6L I-LB-6M I-LBB-7L I-LBB-7M I-LBB-8L I-LBB-8M I-SS-2L I-SS-2M I-SS-3L I-SS-3M I-SS-4L I-SS-4M I-SS-4T I-SS-5M I-SS-6M

Table 2 (continued)

0.335 0.070 0.081 0.112 0.213 0.188 0.260 0.216 0.304 0.213 0.320 0.450 0.182 0.358 0.297 0.054 0.045 0.072 0.053 0.141 0.050 0.210 0.271 0.235 0.354 0.698 1.709 3.202 1.244 3.208 0.059 0.216 0.074 0.074 0.083 0.114 0.223 0.100 0.074

Sm (ppm) 0.072 0.011 0.022 0.016 0.054 0.065 0.066 0.023 0.031 0.026 0.084 0.118 0.039 0.085 0.092 0.020 0.021 0.023 0.013 0.029 0.023 0.057 0.074 0.078 0.087 0.160 0.364 0.649 0.341 0.519 0.018 0.048 0.020 0.015 0.019 0.022 0.016 0.019 0.018

Eu (ppm) 0.160 0.078 0.098 0.096 0.239 0.152 0.264 0.161 0.255 0.163 0.384 0.583 0.231 0.431 0.608 0.081 0.083 0.126 0.093 0.162 0.098 0.320 0.420 0.365 0.463 0.699 1.656 2.936 1.296 2.683 0.076 0.214 0.051 0.096 0.083 0.099 0.093 0.123 0.108

Gd (ppm) 0.036 0.012 0.014 0.008 0.036 0.044 0.051 0.019 0.029 0.026 0.042 0.087 0.041 0.068 0.080 0.020 0.021 0.015 0.012 0.029 0.020 0.055 0.066 0.066 0.073 0.102 0.229 0.362 0.169 0.355 0.008 0.029 0.009 0.014 0.009 0.013 0.018 0.016 0.007

Tb (ppm) 0.267 0.065 0.105 0.080 0.243 0.251 0.219 0.079 0.167 0.142 0.279 0.511 0.207 0.366 0.462 0.126 0.198 0.131 0.126 0.205 0.198 0.409 0.382 0.361 0.481 0.609 1.436 2.057 1.035 1.996 0.069 0.208 0.065 0.115 0.070 0.148 0.075 0.131 0.075

Dy (ppm) 0.085 0.018 0.022 0.020 0.054 0.055 0.054 0.014 0.031 0.027 0.051 0.113 0.059 0.096 0.113 0.039 0.045 0.042 0.029 0.076 0.043 0.090 0.102 0.091 0.106 0.120 0.302 0.408 0.220 0.352 0.015 0.049 0.016 0.033 0.016 0.025 0.022 0.043 0.021

Ho (ppm) 0.192 0.058 0.069 0.065 0.126 0.154 0.141 0.061 0.078 0.075 0.133 0.302 0.147 0.291 0.256 0.107 0.155 0.091 0.085 0.209 0.144 0.248 0.290 0.307 0.298 0.363 0.866 1.146 0.648 0.973 0.042 0.146 0.044 0.098 0.041 0.078 0.054 0.108 0.055

Er (ppm) 0.031 0.008 0.012 0.011 0.017 0.014 0.021 0.006 0.013 0.009 0.015 0.039 0.021 0.041 0.040 0.016 0.015 0.013 0.014 0.022 0.020 0.041 0.037 0.031 0.038 0.048 0.133 0.161 0.093 0.126 0.005 0.021 0.008 0.017 0.007 0.012 0.013 0.016 0.010

Tm (ppm) 0.307 0.051 0.090 0.061 0.099 0.150 0.115 0.040 0.076 0.042 0.102 0.237 0.118 0.199 0.194 0.100 0.100 0.087 0.083 0.240 0.105 0.213 0.177 0.173 0.240 0.344 0.759 0.972 0.566 0.772 0.039 0.128 0.053 0.086 0.041 0.072 0.038 0.108 0.049

Yb (ppm) 0.039 0.007 0.013 0.010 0.017 0.015 0.020 0.006 0.011 0.009 0.013 0.031 0.025 0.034 0.029 0.016 0.013 0.013 0.013 0.026 0.020 0.038 0.038 0.033 0.034 0.046 0.105 0.122 0.079 0.133 0.008 0.019 0.006 0.017 0.007 0.009 0.009 0.016 0.009

Lu (ppm)

C. Su, et al.

Marine and Petroleum Geology 111 (2020) 765–785

777

Marine and Petroleum Geology 111 (2020) 765–785

C. Su, et al.

Table 3 A summary of the elements and oxides, Al2O3, TiO2, Th, and Zr, and the total REE concentrations (ΣREE) of the LMAs from interval B at Laibin section and other sections. Their contents have been used to assess the influence of terrigenous material on carbonates (Kamber et al., 2004; Della Porta et al., 2015; Li et al., 2019). Other sections include interval A in the Laibin, Shangsi, Hongyuan, Jiangkou, Lengshuixi, Laohuangxuan, Guangcan 2, and Lai 1 sections. SD - standard deviation. The element concentrations were acquired using XRF (Al2O3 and TiO2), ICP-MS (Th, Zr, and REEs), and in situ LA-ICP-MS (Al, Ti, Th, Zr, and REEs). The Al and Ti data obtained by the LA-ICP-MS were normalized to the oxides for calculation purposes. Element (bulk rock samples) XRF + ICP-MS

LMAs from Interval B of Laibin (n = 4)

LMAs from other sections (n = 56)

marl from Interval B of Laibin (n = 2)

limestone from Interval B of Laibin (n = 2)

marl from other sections (n = 29)

limestone from other sections (n = 27)

mean

SD

mean

SD

mean

SD

mean

SD

mean

SD

mean

SD

Al2O3 (wt%) TiO2 (wt%) Th (ppm) Zr (ppm) ΣREE (ppm)

5.15 0.188 5.53 100.1 79.9

2.48 0.091 2.51 49.2 21.8

0.36 0.021 0.52 6.5 7.2

0.35 0.015 0.66 6.2 5.7

7.27 0.267 7.68 140.5 98.8

0.68 0.028 0.68 21.3 3.6

3.04 0.109 3.39 59.8 61.1

1.70 0.059 1.71 33.6 15.0

0.55 0.030 0.77 9.4 9.0

0.38 0.017 0.81 6.9 6.6

0.15 0.012 0.25 3.4 5.3

0.08 0.004 0.25 3.1 3.6

Element (thin-sections) LA-ICP-MS

LMAs from Interval B of Laibin (n = 40) mean SD

LMAs from other sections (n = 315) mean SD

marl from Interval B of Laibin (n = 15) mean SD

limestone from Interval B of Laibin (n = 25) mean SD

marl from other sections (n = 166) mean SD

limestone from other sections (n = 114) mean SD

Al2O3 (wt%) TiO2 (wt%) Th (ppm) Zr (ppm) ΣREE (ppm)

2.82 0.100 3.09 27.7 53.8

0.23 0.010 0.24 2.4 5.7

3.78 0.122 4.90 35.1 74.7

2.24 0.086 2.01 23.3 41.3

0.33 0.016 0.34 3.5 6.9

0.11 0.004 0.12 1.0 4.1

1.68 0.066 2.65 19.1 32.5

0.34 0.024 0.31 4.1 4.4

1.17 0.065 3.09 15.4 44.6

1.68 0.062 1.55 19.8 8.7

0.42 0.031 0.35 5.2 5.0

0.17 0.007 0.17 1.3 3.1

microscopic observations, it is generally concluded that the bioclasts in the marl beds were forcefully deformed and oriented (Figs. 3G and 4D, K, and N). In contrast, the fossils in the limestones retain their original shapes and exhibit relatively scattered distributions (Figs. 3F and 4B and M). In addition, fossils with a primary aragonite mineralogy (e.g., gastropods and some bivalves) are common in the limestone beds but are too rare to be recognized in the marl beds (Figs. 3F and 4A). This phenomenon has been explained as the result of selective dissolution, which facilitates differential compaction during early diagenesis (Munnecke and Samtleben, 1996; Melim et al., 2002; Wheeley et al., 2008). This differs from the indiscriminate pressure solution on all of the components that occur in a deep-burial diagenetic setting (Ricken, 1986, 1987). All of the carbonate-dominant LMAs analyzed in this study have homologous features of fossil preservation (Fig. 4B, D, M, and N), implying that they have a similar origin. The widespread distribution of sepiolite (clay mineral; shown as Mg4Si6O15(OH)2·6H2O) indicates the role of early diagenesis in the genesis of the marl beds in the LMAs. Sepiolite (0%–63%) and talc (0%–20%) are the two most dominant clays in the marls in the LMAs from the Shangsi, Hongyuan, Jiangkou, and Laohuangxuan sections (Table 1). The talc is mainly from the transformation of sepiolite (Chen et al., 1985; Yang et al., 1988; Yan et al., 2000, 2005) (Fig. 4P). In some of the sections (e.g., the Shangsi section), extensive sepiolite development in the marl beds leads to the formation of limestone-sepiolite (shale) rhythms (Figs. 3D and 4J and O). We agree that the Permian sepiolite was formed by an early diagenetic process in an environment with sufficient sources of Mg (seawater and the dissolution of high-Mg calcite minerals) and Si (probably silica-secreting organisms) for the precipitation of sepiolite in pore-water conditions (Yan and Carlson, 2003; Yan et al., 2005). The results of the petrographic and LA-ICP-MS analyses also suggest a very limited contribution of terrigenous material to the sepiolite, based on the rarity of the exogenetic siliciclastic minerals and the extremely low concentrations of the siliciclastic-sensitive elements (e.g., Al and Zr; Fig. 6A–D). Furthermore, both the limestone and marl beds (and some transitional layers) hvae similar shale normalized REE + Y patterns (Fig. 7C–H). In particular, the LMAs in interval A of the Laibin section have typical modern seawater-like REE + Y patterns characterized by negative Ce and positive La anomalies, enriched HREEs relative to LREEs, and superchondritic Y/Ho ratios (> 40) (see Section 5.2.2; Nozaki, 2001). These patterns resemble those of modern nearly pure carbonate sediments (microbialites and ooids)

analyses show much higher Al2O3, Zr, and ΣREE contents within the marl and limestone beds of the LMAs in interval B of the Laibin section than in the other sites (Table 3 and Fig. 6), implying a more intense terrigenous input during the formation of the primary LMA deposits, especially in the marls (Westphal et al., 2004; Frimmel, 2009). Fourth, REE + Y characteristics in interval B indicate two different sources for the coupled limestone and marl beds (Fig. 7A and B). This is because the consistent primary components of the limestone and the neighboring marl would produce similar REE + Y features in the preserved LMAs if they experienced similar diagenetic processes. However, the original carbonate REE + Y signatures have been mostly overprinted by the terrigenous components, and thus, they are characterized by flattened shale-normalized REE + Y patterns (Fig. 7A and B) and inconsistent ΣREE concentrations (Table 3), which were most likely caused by differential siliciclastic input. Based on the above described four reasons, we suggest that the non-carbonate components of the limestone and marl beds are varied and the heterogeneity of the terrigenous supply to the two kinds of beds may have contributed to the generation of the LMAs in interval B of the Qixia Formation at Laibin section. 5.2. Permian diagenetic LMAs: A universal phenomenon in relatively “pure carbonate” depositional settings? 5.2.1. Evidence of carbonate diagenesis in the formation of LMAs Carbonate diagenesis is regarded as a key factor in the generation of LMAs (Hallam, 1986; Ricken, 1986, 1987; Raiswell, 1987; Munnecke and Samtleben, 1996; Munnecke et al., 1997, 2008; Reinhardt et al., 2000; Holmes et al., 2004; Westphal et al., 2004, Westphal et al., 2008a; 2010, 2015; Bádenas et al., 2009; Beltran et al., 2009; Li et al., 2018). Evidence from fossil preservation (Munnecke and Samtleben, 1996; Melim et al., 2002; Zhan et al., 2016), mineralogical composition (Munnecke and Westphal, 2005; Westphal et al., 2010; Li et al., 2018), and major element geochemistry (Westphal et al., 2004; Biernacka et al., 2005; Amberg et al., 2016) has been used to investigate the influence of differential diagenesis on the LMAs. In this study, the characteristics of eight sections (except for interval B of the Laibin section), which formed in various environments with rare to minor terrigenous supplies, are thoroughly evaluated in order to determine and analyze the reliable diagenetic characteristics of the Permian LMAs. The orientation of the skeletal fragments in the marl beds implies the occurrence of differential compaction. Based on both field and 778

Marine and Petroleum Geology 111 (2020) 765–785

C. Su, et al.

Fig. 6. (A) Plot of the Y/Ho mass ratio vs. Al2O3 (wt%) from bulk rock samples (method from Tostevin et al., 2016). (B) Plot of Y/Ho vs. Al2O3 of the in situ LA-ICPMS data. (C) and (D) Plots of Y/Ho vs. Zr of the solution-based ICP-MS and LA-ICP-MS data, respectively. (E) and (F) Plots of Y/Ho vs. ΣREE of the solution-based ICPMS and LA-ICP-MS data, respectively.

(Webb and Kamber, 2000; Li et al., 2019). The additional supply of siliciclastic components to the marl beds would significantly elevate the shale-normalized LREE patterns (Della Porta et al., 2015; Li et al., 2019) compared to those of the coupled limestone beds. However, this elevation is not seen in the studied sections containing relatively pure LMAs (Fig. 7C–H). In addition, in South China, previous exposed areas were drowned during the Roadian (Wang and Jin, 2000) (Fig. 1), so their ability to supply siliciclastic material to carbonate depositional area (LMAs) was weak. Therefore, the relatively high volumes of sepiolite and talc in the LMAs do not reflect the true depositional differences between the limestone and marl beds.

key role of carbonate diagenesis in the genesis of the Permian LMAs. Munnecke et al. (1997) and Westphal et al. (2008a, 2015) believe the existence of LMA completely controlled by carbonate diagenesis, but such instances are rare, and the verification method is not convincing (e.g., Eldrett et al., 2015a; Amberg et al., 2016). In this study, REE + Y compositions are used to investigate the causes of the geochemical signatures of the limestone and marl beds in the LMAs. This is because the behavior of REE + Y in carbonate is more conservative than the behavior of other metal elements, and their characteristics are largely preserved during diagenetic processes when the interference of exogenetic particles/fluids is negligible (Nothdurft et al., 2004; Webb et al., 2009; Hood et al., 2018; Li et al., 2019). The REE + Y signatures of of the limestone and marl beds are consistent in interval A at Laibin section, implying that they had identical original compositions. The limestone and marl beds in interval A at Laibin section (I-LBA-5) exhibit uniform REE + Y patterns with the following characteristics. (1) The limestone, marl, and transitional beds

5.2.2. Do purely diagenetic LMAs exist? Although evidence of carbonate diagenesis is present in at all of the siliciclastic-insufficient sections analyzed in this study (excluding interval B at Laibin section), it is still unclear whether entirely diagenetic LMAs exist. If present, they would provide solid evidence to support the 779

Marine and Petroleum Geology 111 (2020) 765–785

C. Su, et al.

Fig. 7. PAAS-normalized REE + Y patterns of the LMA successions. (A) and (B) Interval B in the Laibin section. (C) and (D) Interval A in the Laibin section. (E) Hongyuan section. (F) Well Guangcan 2. (G) and (H) Shangsi section. Detailed sampling positions are shown in Fig. 2; and original REE + Y data are listed in the Table 2.

have overlapping shale-normalized REE + Y patterns. (2) All three components have diagnostic seawater-like patterns with positive La and negative Ce anomalies ((La/La*)SN = 1.64 ± 0.17, (Ce/ Ce*)SN = 0.73 ± 0.05; method from Lawrence et al., 2006), superchondritic Y/Ho (47 ± 4), and significant LREE depletion relative to the HREEs ((Pr/Yb)SN = 0.39 ± 0.09). (3) They have high volumes of CaCO3 and are almost free of clays (Figs. 5 and 7C; Table 1). Terrigenous quartz grains have never been microscopically identified in marl beds (Fig. 4B–E, K, and L), except for some fossils replaced by quartz (inferred to be the result of diagenesis; Fig. 4M and 8). Therefore, this is a good example of the role of carbonate diagenesis in the formation of LMAs originating from two pure limestones. Another example from the Hongyuan section supports the existence of diagenetic LMAs with the slight contamination of detrital materials.

The limestone and marl beds of the Hongyuan LMAs have overlapping REE signatures but have less depleted of LREE patterns compared to those of interval A at Laibin section (Fig. 7E; (Pr/Yb)SN = 0.80 ± 0.03, n = 3; note that the original YbSN value of the marl bed is anomalous and has been replaced by the mean value of the adjacent TmSN and LuSN values). The volumes of siliciclastics added to the two beds must be very similar; otherwise, the marl bed would have a heavier REE signature and a flatter LREE pattern. However, only a few siliciclastic impurities have been mixed with the carbonate sediments of the Hongyuan section. This is because siliciclastic components generally contain extremely high REE concentrations (e.g., ΣREEsilts = 100–260 ppm and ΣREEclays = 110–300 ppm), which would significantly change the carbonate REE concentrations and patterns (Bayon et al., 2015) (Fig. 7A and B). In addition, the marked differential preservation of fossils is 780

Marine and Petroleum Geology 111 (2020) 765–785

C. Su, et al.

Fig. 8. Brachiopod fragment that has been replaced by quartz at in a marl bed in an LMA succession. Interval A in the Laibin section. (A) Plane-polarized light. (B) Cross-polarized light.

seen in the marl beds in the LMAs from the Hongyuan section. In summary, a small amount of terrigenous input occurred during carbonate sedimentation, but its contribution to the genesis of the compacted (dissolved) marl beds is weak compared to diagenetic alteration.

of generated sulfides diffuse and reoxidize to sulfate when the overlying dissolved oxygen is sufficient (Visscher and Stolz, 2005). The relatively reduced pore-water eventually finalizes the stabilization of the aragonite and high-Mg calcite. In general, active aerobic sulfide oxidation during early diagenesis may promote to the formation of diagenetic LMAs. In addition, high volumes of terrigenous input may provide the marls with abundant Fe oxides (Fig. 9), which would generate an alkaline microenvironment and promot cementation during Fe reduction in an early diagenetic phrase (e.g., Best et al., 2007, 2008). This process may delay the transformation of the metastable aragonite and high-Mg calcite and can be inferred as the reason for the negligible diagenetic alteration of the sedimentary LMAs.

5.2.3. The genesis of diagenetic LMAs in the Permian strata of South China: A hypothesis Here, an updated model is proposed to shed light on the genesis of diagenetic LMAs in the Permian strata of South China based on previous understandings of their development (Munnecke and Samtleben, 1996; Westphal et al., 2000, 2010, 2015; Böhm et al., 2003; Yan et al., 2005). High contents of organic matter and sulfate-enriched pore-water fluids are two potential reasons for the formation of diagenetic LMAs. The occurrences of large quantities of calcareous algae and other fossils in the Permian sedimentary record provide sufficient biomass and suggest high primary productivity rates (Fig. 4D, K, and N). After this, shallowly buried bioclasts and early cements with primary aragonitic and highMg calcite mineralogies would transform into stable low-Mg calcites during early diagenesis (Land, 1967), resulting in the stepwise loss of organic matter when the pore-water is unsaturated in relation to aragonite (dominate) and high-Mg calcite. Because aragonite dissolution requires decreased pH conditions, sulfide oxidation below the oxic zone is probable (Aller, 1982; Boudreau and Canfield, 1993). The interval between oxic respiration and sulfate reduction is a favorable environment for aragonite dissolution when the Fe and Mn reduction zones are very narrow (Visscher et al., 1998; Petrash et al., 2017). The state of atmospheric oxygen in the Permian was probably highest during the Phanerozoic (Berner, 2009), and thus, the oxic zone in the pore-water may generally be deeper than during other periods. In addition, Permian seawater had very high concentrations of sulfate (~20 mM) (Lowenstein et al., 2005). Such high values can be inherited by porewater and can facilitate to sulfate reduction in anoxic pore-water. A lot

5.3. New understandings of the Permian calcareous LMA successions 5.3.1. The application of LMAs in cyclostratigraphy Orbitally forced Milankovitch cycles can provide useful information on variations in climate, environment, and sea-level (Fischer and Bottjer, 1991; Elder et al., 1994; Dinarès-Turell et al., 2003, 2010; Strasser et al., 2006; Berrocoso et al., 2013; Ma et al., 2017). A series of alternating carbonate and siliciclastic successions (mainly marine and lake deposits) have been extensively used in past glacio-eustatic and astrochronological studies (Hilgen, 1987; King, 1990; Pittet et al., 2000; Montgomery et al., 2001; Kuhn and Diekmann, 2002; Roth and Reijmer, 2005). Previous studies warn that due to differential diagenesis, under certain conditions, the original depositional and geochemical characteristics of carbonate-dominant LMAs may be altered, e.g., the thickness, magnetic susceptibility, mineralogy, and element preservation (e.g., Böhm et al., 2003; Westphal, 2006; Westphal et al., 2008a, 2008b; 2010). However, several studies stress the primary compositional differences between the limestone and marl beds and believe that the influence of carbonate diagenesis is insignificant Fig. 9. Plots of Fe vs. Zr showing the influence of Fe on the preservation of the LMAs. (A) The good correlation between Fe and Zr suggests the Fe is principally from detrital input. (B) The rare to minor Fe characteristics are generally preserved in the diagenetic LMAs (Ⅰ). Significantly different Fe concentrations in the limestone and marl beds suggest different compositions for the primary LMAs (Ⅱ). Note that not all LMAs from the same section have the same origins. The Fe data in (A) are from the XRF and ICP-OES analyses.

781

Marine and Petroleum Geology 111 (2020) 765–785

C. Su, et al.

Fig. 10. Plot of Al2O3 vs. TiO2 of the Permian LMAs from South China. Data from bulk rock (coupled limestone and marl beds) XRF and ICP-OES analyses. Note that not all LMAs in interval A at Laibin section are interpreted to have purely diagenetic origins.

(Bellanca et al., 1996; Huang and Baus, 1999; Eldrett et al., 2015a; Elderbak and Leckie, 2016). The fully diagenetic LMAs identified in this study provide compelling evidence of the role of diagenesis in the formation of LMAs originating from “pure carbonate” successions (Fig. 7C). Carbonate-dominant LMAs containing a small amount of detritus also exhibit the diagnostic characteristics of differential diagenesis (e.g., the Hongyuan section; Fig. 7E). Thus, both the pure and slightly-contaminated shallow-water carbonate LMA successions in the Permian strata of South China cannot be convincingly applied in the study of cyclostratigraphy.

from diagenetic alteration. Second, we confirm the development of purely diagenetic LMAs derived from two pristine limestone beds without the participation of terrigenous material. Third, rare to minor admixtures of terrigenous input likely have a weak effect to diagenetic alteration, but a larger supply of terrigenous material alters or completely overprints the diagenetic signature of the marl bed in the LMA. Fourth, an updated model was provided to explain the widespread distribution of LMAs with a limited terrigenous supply in South China during the Permian. This suggests that the formation of diagenetic LMAs depends on the active metabolic processes involved in the decomposition of (sufficient) organic matter during early diagenesis, especially aerobic sulfide oxidation, and suitable palaeoceanographic conditions (high levels of dissolved oxygen and sulfate). The specific processes in this model need further investigation.

5.3.2. Limitation of Al2O3 vs. TiO2 in identifying the signatures of diagenetic alteration Both the Al2O3 and TiO2 contents originate from the terrigenous supply, and their crossplot is generally used to trace the input of detrital components into LMAs (Westphal et al., 2004, 2006; Westphal et al., 2008a, 2008b; 2010; Biernacka et al., 2005; Munnecke et al., 2008; Eldrett et al., 2015a; Amberg et al., 2016). Inconsistent slopes of the regression lines of Al2O3 vs. TiO2 acquired from the limestone and marl beds indicate that they have varied terrigenous sources, suggesting a primary depositional difference within the LMA (Westphal et al., 2004, Westphal et al., 2008a). However, the significance of the nearly uniform slopes of the regression lines of Al2O3 vs. TiO2 of coupled limestone and marl beds is uncertain. They may indicate either the same detrital source, or they may be the result of diagenesis of the limestone precursor (Westphal et al., 2010; Amberg et al., 2016). In this study, the different LMA components from all eight sections exhibit a very good linearly-correlated of the regression lines for Al2O3 vs. TiO2, although they have different depositional settings and contain rare to abundant siliciclastic detritus (Fig. 10; n = 60, R2 = 0.99). For example, the diagenetic and sedimentary LMAs from interval A (I-LBA-5) and interval B (I-LBA-7 and -8) of the Laibin section, respectively, have nearly consistent slopes of the regression lines of Al2O3 vs. TiO2 (Fig. 10). Thus, we suggest that Al2O3 vs. TiO2 is not capable of determining the genesis of the LMAs in the Permian strata of South China.

Acknowledgments We thank editor Xiaomin Zhu and three anonymous reviewers for their constructive comments that were helpful to improve the quality of the manuscript. We thank Feifan Lu and Xinyu Zhang for their assistance in the field. This study was jointly supported by the National Natural Science Foundation of China (Nos. 41872119 and 41502115), the National Science and Technology Major Project (No. 2016ZX05004002-001), the Key Laboratory of Carbonate Reservoirs, CNPC (Nos. RIPED-HZDZY-2018-JS-198 and 2018D-5006-35), and the IAS Post-Doctoral Research Grants (Spring Session, 2017). Appendix A. Supplementary data Supplementary data to this article can be found online at https:// doi.org/10.1016/j.marpetgeo.2019.08.033. References Aller, R.C., 1982. Carbonate dissolution in nearshore terrigenous muds: the role of physical and biological reworking. J. Geol. 90, 79–95. Amberg, C.E., Collart, T., Salenbien, W., Egger, L.M., Munnecke, A., Nielsen, A.T., Monnet, C., Hammer, O., Vandenbroucke, T.R., 2016. The nature of Ordovician limestone–marl alternations in the Oslo–Asker District (Norway): witnesses of primary glacio–eustasy or diagenetic rhythms? Sci. Rep. 6, 1–13. Bádenas, B., Aurell, M., García–Ramos, J.C., González, B., Piñuela, L., 2009. Sedimentary vs. diagenetic control on rhythmic calcareous successions (Pliensbachian of Asturias, Spain). Terra. Nova 21, 162–170. Bádenas, B., Aurell, M., Armendáriz, M., Rosales, I., García–Ramos, J.C., Piñuela, L.,

6. Conclusions In this study, we provided insights into the genesis of the LMAs in the Permian of South China based on petrographic, mineralogical, and element geochemical evidence. First, there are two different kinds of LMAs, which originated from primary sedimentary differences and

782

Marine and Petroleum Geology 111 (2020) 765–785

C. Su, et al. 2012. Sedimentary and chemostratigraphic record of climatic cycles in Lower Pliensbachian marl–limestone platform successions of Asturias (North Spain). Sediment. Geol. 281, 119–138. Bauer, A.J., Burte, E.P., Hornak, J., Rolls, E.T., Wade, D., Ghafouri–Shiraz, H., Aruga, T., Elrick, M., Hinnov, L.A., 1996. Millennial–scale climate origins for stratification in Cambrian and Devonian deep–water rhythmites, western USA. Palaeogeogr. Palaeoclimatol. Palaeoecol. 123, 353–372. Bayon, G., Toucanne, S., Skonieczny, C., André, L., Bermell, S., Cheron, S., Dennielou, B., Etoubleau, J., Freslon, N., Gauchery, T., Germain, Y., Jorry, S.J., Ménot, G., Monin, L., Ponzevera, E., Rouget, M.L., Tachikawa, K., Barrat, J.A., 2015. Rare earth elements and neodymium isotopes in world river sediments revisited. Geochem. Cosmochim. Acta 170, 17–38. Bellanca, A., Claps, M., Erba, E., Masetti, D., Neri, R., Silva, I.P., Venezia, F., 1996. Orbitally induced limestone/marlstone rhythms in the Albian—cenomanian Cismon section (Venetian region, northern Italy): sedimentology, calcareous and siliceous plankton distribution, elemental and isotope geochemistry. Palaeogeogr. Palaeoclimatol. Palaeoecol. 126, 227–260. Bellanca, A., Masetti, D., Neri, R., 1997. Rare earth elements in limestone/marlstone couplets from the Albian–Cenomanian Cismon section (Venetian region, northern Italy): assessing REE sensitivity to environmental changes. Chem. Geol. 141, 141–152. Beltran, C., Rafélis, M.D., Person, A., Stalport, F., Renard, M., 2009. Multiproxy approach for determination of nature and origin of carbonate micro–particles so–called “micarb” in pelagic sediments. Sediment. Geol. 213, 64–76. Berner, R.A., 2009. Phanerozoic atmospheric oxygen: new results using the GEOCARBSULF model. Am. J. Sci. 309, 603–606. Berrocoso, Á., Elorza, J., Macleod, K.G., 2013. Proximate environmental forcing in fine–scale geochemical records of calcareous couplets (Upper Cretaceous and Palaeocene of the Basque–Cantabrian Basin, eastern North Atlantic). Sediment. Geol. 284–285, 76–90. Best, M.M.R., 2008. Contrast in preservation of bivalve death assemblages in siliciclastic and carbonate tropical shelf settings. Palaios 23, 796–809. Best, M.M.R., Ku, T.C.W., Kidwell, S.M., Walter, L.M., 2007. Carbonate preservation in shallow marine environments: unexpected role of tropical siliciclastics. J. Geol. 115, 437–456. Biernacka, J., Borysiuk, K., Raczyński, P., 2005. Zechstein (Ca1) limestone–marl alternations from the North–Sudetic Basin, Poland: depositional or diagenetic rhythms? Geol. Q. 49, 1–14. Böhm, F., Westphal, H., Bornholdt, S., 2003. Required but disguised: environmental signals in limestone–marl alternations. Palaeogeogr. Palaeoclimatol. Palaeoecol. 189, 161–178. Boudreau, B.P., Canfield, D.E., 1993. A comparison of closed- and open-system models for porewater pH and calcite-saturation state. Geochem. Cosmochim. Acta 57, 317–334. Boulila, S., de Rafélis, M., Hinnov, L.A., Gardin, S., Galbrun, B., Collin, P.–Y., 2010. Orbitally forced climate and sea–level changes in the Paleoceanic Tethyan domain (marl–limestone alternations, Lower Kimmeridgian, SE France). Palaeogeogr. Palaeoclimatol. Palaeoecol. 292, 57–70. Chamley, H., 1998. Clay mineral sedimentation in the Ocean. In: Paquet, H., Clauer, N. (Eds.), Soils and Sediments (Mineralogy and Geochemistry). Springer-Verlag, Berlin, pp. 269–302. Chen, Y., Wang, P., Ren, L., 1985. Study on the transformation of sepiolite to talc during diagenesis. Chin. Sci. Bull. 4, 284–287 (in Chinese). Chen, L., Liu, Y.S., Hu, Z.C., Gao, S., Zong, K.Q., Chen, H.H., 2011. Accurate determinations of fifty–four major and trace elements in carbonate by LA–ICP–MS using normalization strategy of bulk components as 100%. Chem. Geol. 284, 283–295. Coimbra, R., Immenhauser, A., Olóriz, F., Rodríguez–Galiano, V., Chica–Olmo, M., Pufahl, P., 2015. New insights into geochemical behaviour in ancient marine carbonates (Upper Jurassic Ammonitico Rosso): novel proxies for interpreting sea–level dynamics and palaeoceanography. Sedimentology 62, 266–302. Courtinat, B., 1993. The significance of palynofacies fluctuations in the greenhorn formation (Cenomanian–Turonian) of the western interior basin, USA. Mar. Micropaleontol. 21, 249–257. Della Porta, G., Webb, G.E., McDonald, I., 2015. REE patterns of microbial carbonate and cements from Sinemurian (Lower Jurassic) siliceous sponge mounds (Djebel Bou Dahar, High Atlas, Morocco). Chem. Geol. 400, 65–86. Dinarès–Turell, J., 2010. High–resolution intra– and interbasinal correlation of the danian–selandian transition (early paleocene): the bjala section (Bulgaria) and the selandian GSSP at zumaia (Spain). Palaeogeogr. Palaeoclimatol. Palaeoecol. 297, 511–533. Dinarès–Turell, J., Baceta, J.I., Pujalte, V., Orue–Etxebarria, X., Bernaola, G., Lorito, S., 2003. Untangling the Palaeocene climatic rhythm: an astronomically calibrated Early Palaeocene magnetostratigraphy and biostratigraphy at Zumaia (Basque basin, northern Spain). Earth Planet. Sci. Lett. 216, 483–500. Einsele, G., Ricken, W., 1991. Limestone–marl alternation – an overview. In: Einsele, G., Ricken, W., Seilacher, A. (Eds.), Orbital Forcing Timescales and Cyclostratigraphy. vol. 85. Geological Society, London, Special Publication, pp. 23–47. Einsele, G., Seilacher, A., 1982. Cyclic and Event Stratification. Springer–Verlag, Berlin Heidelberg New York. Einsele, G., Ricken, W., Seilacher, A., 1991. Cycles and Events in Stratigraphy. Springer–Verlag, Berlin. Elder, W.P., Gustason, E.R., Sageman, B.B., 1994. Correlation of basinal carbonate cycles to nearshore parasequences in the Late Cretaceous Greenhorn seaway, Western

Interior U.S.A. Geol. Soc. Am. Bull. 106, 892–902. Elderbak, K., Leckie, R.M., 2016. Paleocirculation and foraminiferal assemblages of the Cenomanian–Turonian Bridge Creek Limestone bedding couplets: Productivity vs. dilution during OAE2. Cretac. Res. 60, 52–77. Eldrett, J.S., Ma, C., Bergman, S.C., Ozkan, A., Minisini, D., Lutz, B., Jackett, S.–J., Macaulay, C., Kelly, A.E., 2015a. Origin of limestone–marlstone cycles: Astronomic forcing of organic–rich sedimentary rocks from the Cenomanian to early Coniacian of the Cretaceous Western Interior Seaway, USA. Earth Planet. Sci. Lett. 423, 98–113. Eldrett, J.S., Ma, C., Bergman, S.C., Lutz, B., Gregory, F.J., Dodsworth, P., Phipps, M., Hardas, P., Minisini, D., Ozkan, A., 2015b. An astronomically calibrated stratigraphy of the Cenomanian, Turonian and earliest Coniacian from the Cretaceous Western Interior Seaway, USA: Implications for global chronostratigraphy. Cretac. Res. 56, 316–344. Elrick, M., Hinnov, L.A., 2007. Millennial–scale paleoclimate cycles recorded in widespread Palaeozoic deeper water rhythmites of North America. Palaeogeogr. Palaeoclimatol. Palaeoecol. 243, 348–372. Eslinger, E., Pevear, D., 1988. Clay Minerals for Petroleum Geologists and Engineers. SEPM Short Course Notes, pp. 22. Fang, Q., Jing, X.C., Deng, S.H., Wang, X.L., 2012. Roadian–Wuchiapingian conodont biostratigraphy at the Shangsi section, northern Sichuan. J. Stratigr. 36, 692–699 (in Chinese with English abstract). Feng, Z., Yang, Y., Jin, Z., He, Y.B., Wu, S.H., 1996. Lithofacies paleogeography of the Permian of South China. Acta Sedimentol. Sin. 14, 1–11 (in Chinese with English abstract). Fischer, A.G., Bottjer, D.J., 1991. Orbital forcing and sedimentary sequences. J. Sediment. Petrol. 61, 1063–1069. Frimmel, H.E., 2009. Trace element distribution in Neoproterozoic carbonates as palaeoenvironmental indicator. Chem. Geol. 258, 338–353. Gardulski, A.F., Mullins, H.T., Weiterman, S., 1990. Carbonate mineral cycles generated by foraminiferal and pteropod response to Pleistocene climate: west Florida ramp slope. Sedimentology 37, 727–743. Haley, B.A., Klinkhammer, G.P., McManus, J., 2004. Rare earth elements in pore waters of marine sediments. Geochem. Cosmochim. Acta 68, 1265–1279. Hallam, A., 1986. Origin of minor limestone–shale cycles: Climatically induced or diagenetic? Geology 14, 609–612. Haq, B.U., Schutter, S.R., 2008. A chronology of Paleozoic sea–level changes. Science 322, 64–68. Hattin, D.E., 1971. Widespread, synchronously deposited, burrow–mottled limestone beds in Greenhorn Limestone (Upper Cretaceous) of Kansas and central Colorado. AAPG Bull. 55, 412–431. He, B., Xu, Y.G., Wang, Y.M., Xiao, L., 2005. Nature of the Dongwu movement and its temporal and spatial evolution. Earth Sci. J. China Univ. Geosci. 30, 89–96 (in Chinese with English abstract). Hilgen, F.J., 1987. Sedimentary rhythms and high–resolution chronostratigraphic correlations in the Mediterranean Pliocene. Newsl. Stratigr. 17, 109–127. Holmes, M.A., Watkins, D.K., Norris, R.D., 2004. Paleocene cyclic sedimentation in the western North Atlantic, ODP Site 1051, Blake Nose. Mar. Geol. 209, 31–43. Hood, A.V.S., Planavsky, N.J., Wallace, M.W., Wang, X.L., 2018. The effects of diagenesis on geochemical paleoredox proxies in sedimentary carbonates. Geochem. Cosmochim. Acta 232, 265–287. Hooper, P.R., 1969. The preparation of fused samples in X–Ray fluorescence analysis. Mineral. Mag. 37, 409–413. Huang, J., 1985. The sedimentary environment pattern and its tectonic control of Late Permian Age in Nanpanjiang River area. J. Southwest Pet. Inst. 1–10 (in Chinese with English abstract). Huang, W.H., Baus, W.M., 1999. The geochemical feature’s comparison between the typical rhythmic bedding Iimestone and marl of upper Jurassic Malm in southern Germany. Acta Sedimentol. Sin. 17, 633–637 (in Chinese with English abstract). Husson, D., Thibault, N., Galbrun, B., Gardin, S., Minoletti, F., Sageman, B., Huret, E., 2014. Lower Maastrichtian cyclostratigraphy of the Bidart section (Basque Country, SW France): A remarkable record of precessional forcing. Palaeogeogr. Palaeoclimatol. Palaeoecol. 395, 176–197. International Commission on Stratigraphy, 2017. International Chronostratigraphic Chart. Kamber, B.S., Bolhar, R., Webb, G.E., 2004. Geochemistry of late Archaean stromatolites from Zimbabwe: evidence for microbial life in restricted epicontinental seas. Precambrian Res. 132, 379–399. King Jr., D.T., 1990. Upper Cretaceous marl–limestone sequences of Alabama: Possible products of sea–level change, not climate forcing. Geology 18, 19. Koch, J.T., Frank, T.D., 2011. The Pennsylvanian–Permian transition in the low–latitude carbonate record and the onset of major Gondwanan glaciation. Palaeogeogr. Palaeoclimatol. Palaeoecol. 308, 362–372. Kuhn, G., Diekmann, B., 2002. Late quaternary variability of the ocean circulation in the southeastern South Atlantic inferred from terrigenous sediment record of a drift deposit in the southern Cape Basin (ODP Site 1089). Palaeogeogr. Palaeoclimatol. Palaeoecol. 182, 287–303. Land, L.S., 1967. Diagenesis of skeletal carbonates. J. Sediment. Res. 37, 914–930. Lathuilière, B., Bartier, D., Bonnemaison, M., Boullier, A., Carpentier, C., Elie, M., Gaillard, C., Gauthier–Lafaye, F., Grosheny, D., Hantzpergue, P., 2015. Deciphering the history of climate and sea level in the Kimmeridgian deposits of Bure (eastern Paris Basin). Palaeogeogr. Palaeoclimatol. Palaeoecol. 433, 20–48. Lawrence, M.G., Greig, A., Collerson, K.D., Kamber, B.S., 2006. Rare earth element and

783

Marine and Petroleum Geology 111 (2020) 765–785

C. Su, et al. yttrium variability in South East Queensland waterways. Aquat. Geochem. 12, 39–72. Li, F., Yan, J., Algeo, T., Wu, X., 2013. Paleoceanographic conditions following the end–Permian mass extinction recorded by giant ooids (Moyang, South China). Glob. Planet. Chang. 105, 102–120. Li, F., Yan, J., Burne, R.V., Chen, Z., Algeo, T.J., Zhang, W., Tian, L., Gan, Y., Liu, K., Xie, S., 2017. Paleo–seawater REE compositions and microbial signatures preserved in laminae of Lower Triassic ooids. Palaeogeogr. Palaeoclimatol. Palaeoecol. 486, 96–107. Li, F., Webb, G.E., Algeo, T.J., Kershaw, S., Lu, C., Oehlert, A.M., Gong, Q., Pourmand, A., Tan, X., 2019. Modern carbonate ooids preserve ambient aqueous REE signatures. Chem. Geol. 509, 163–177. Li, J., Cai, Z., Chen, H., Cong, F., Wang, L., Wei, Q., Luo, Y., 2018. Influence of differential diagenesis on primary depositional signals in limestone–marl alternations: An example from Middle Permian marine successions, South China. Palaeogeogr. Palaeoclimatol. Palaeoecol. 495, 139–151. Liu, Y., Hu, Z., Gao, S., Günther, D., Xu, J., Gao, C., Chen, H., 2008. In situ analysis of major and trace elements of anhydrous minerals by LA–ICP–MS without applying an internal standard. Chem. Geol. 257, 34–43. Liu, X., Yan, J., Xue, W., 2012. Differential diagenesis of limestone–marl alternations. Geol. Rev. 58, 627–635 (in Chinese with English abstract). Liu, C., Lu, G., Zhang, X., Zhou, F., Huang, X., 2014. Sedimentary characteristics and evolutionary stages in Late Paleozoic of Tian’e isolated platform in northwestern Guangxi. Geol. Rev. 60, 55–70 (in Chinese with English abstract). Liu, X., Yan, J., Ma, Z., Xue, W., 2014. Origination of limestone–marl alternations from Qixia Formation of South China. J. China Univ. Geosci. 39, 155–164 (in Chinese with English abstract). Lowenstein, T.K., Timofeeff, M.N., Kovalevych, V.M., Horita, J., 2005. The major-ion composition of Permian seawater. Geochem. Cosmochim. Acta 69, 1701–1719. Luo, J., He, Y., 2010. Origin and characteristics of Permian eyeball–shaped limestones in middle—upper Yangtze region. Geol. Rev. 56, 629–637 (in Chinese with English abstract). Luo, Q., Hou, F., 1990. Study of Permian to Lower Triassic depositional systems in Pingguo area, Guangxi. J. Southwest Pet. Inst. 12, 1–8 (in Chinese with English abstract). Ma, Z., Yan, J., Xie, X., Ruan, X., Li, B., 2008. Depositional and ecological features of Permian oxygen deficient deposits at Shangsi section, northeast Sichuan, China. J. China Univ. Geosci. 19, 488–495. Ma, C., Meyers, S.R., Sageman, B.B., 2017. Theory of chaotic orbital variations confirmed by Cretaceous geological evidence. Nature 542, 468. Melim, L.A., Westphal, H., Swart, P.K., Eberli, G.P., Munnecke, A., 2002. Questioning carbonate diagenetic paradigms: evidence from the Neogene of the Bahamas. Mar. Geol. 1, 27–53. Montgomery, P., Farr, M.R., Franseen, E.K., Goldstein, R.H., 2001. Constraining controls on carbonate sequences with high–resolution chronostratigraphy: Upper Miocene, Cabo de Gata region, SE Spain. Palaeogeogr. Palaeoclimatol. Palaeoecol. 176, 11–45. Munnecke, A., Samtleben, C., 1996. The formation of micritic limestones and the development of limestone–marl alternations in the Silurian of Gotland, Sweden. Facies 34, 159–176. Munnecke, A., Westphal, H., 2005. Variations in primary aragonite, calcite, and clay in fine–grained calcareous rhythmites of Cambrian to Jurassic age— an environmental archive? Facies 51, 592–607. Munnecke, A., Westphal, H., Reijmer, J.J.G., Samtleben, C., 1997. Microspar development during early marine burial diagenesis: a comparison of Pliocene carbonates from the Bahamas with Silurian limestones from Gotland (Sweden). Sedimentology 44, 977–990. Munnecke, A., Westphal, H., Kölbl–Ebert, M., 2008. Diagenesis of plattenkalk: examples from the Solnhofen area (Upper Jurassic, southern Germany). Sedimentology 55, 1931–1946. Nothdurft, L.D., Webb, G.E., Kamber, B.S., 2004. Rare earth element geochemistry of Late Devonian reefal carbonates, Canning Basin, Western Australia: confirmation of a seawater REE proxy in ancient limestones. Geochem. Cosmochim. Acta 68, 263–283. Nozaki, Y., 2001. Rare earth elements and their isotopes in the ocean. In: Steele, J.H., Thorpe, S.A., Turekian, K.K. (Eds.), Encyclopedia of Ocean Sciences. Academic Press, Oxford, pp. 2354–2366. Olesik, J.W., 1991. Elemental analysis using ICP–OES and ICP–MS. Anal. Chem. 63, 187–192. Petrash, D.A., Bialik, O.M., Bontognali, T.R.R., Vasconcelos, C., Roberts, J.A.k, McKenzie, J.A., Konhauser, K.O., 2017. Microbially catalyzed dolomite formation: from nearsurface to burial. Earth Sci. Rev. 171, 558–582. Pittet, B., Strasser, A., Mattioli, E., 2000. Depositional sequences in deep–shelf environments: A response to sea–level changes and shallow–platform carbonate productivity (Oxfordian, Germany and Spain). J. Sediment. Res. 70, 392–407. Planavsky, N., Bekker, A., Rouxel, O.J., Kamber, B., Hofmann, A., Knudsen, A., Lyons, T.W., 2010. Rare Earth Element and yttrium compositions of Archean and Paleoproterozoic Fe formations revisited: New perspectives on the significance and mechanisms of deposition. Geochem. Cosmochim. Acta 74, 6387–6405. Raiswell, R., 1987. Non–steady state microbiological diagenesis and the origin of concretions and nodular limestones. Geol. Soc. London Spec. Publications 36, 41–54. Reinhardt, E.G., Cavazza, W., Patterson, R.T., Blenkinsop, J., 2000. Differential diagenesis of sedimentary components and the implication for strontium isotope analysis of carbonate rocks. Chem. Geol. 164, 331–343. Ricken, W., 1986. Diagenetic Bedding : a Model for Marl–Limestone Alternations. Lecture

Notes in Earth Sciences Berlin Springer Verlag, pp. 6. Ricken, W., 1987. The carbonate compaction law: a new tool. Sedimentology 34, 571–584. Ross, C.A., Ross, J.R.P., 1985. Late Paleozoic depositional sequences are synchronous and worldwide. Geology 13, 194–197. Roth, S., Reijmer, J.J.G., 2005. Holocene millenial to centennial carbonate cyclicity recorded in slope sediments of the Great Bahama Bank and its climatic implications. Sedimentology 52, 161–181. Scotese, C.R., 2014. Atlas of Middle & Late Permian and Triassic Paleogeographic Maps, Maps 43–52 from Volumes 3 & 4 of the PALEOMAP Atlas for ArcGIS, Mollweide Projection. PALEOMAP Project, Evanston. Scotese, C.R., Langford, R.P., 1995. Pangea and the Paleogeography of the Permian. Springer Berlin Heidelberg. Seibold, E., 1952. Chemische Untersuchungen zur Kalk–Mergel–Sedimentation. Geol. Rundsch. 40 284–284. Sha, Q., Wu, W., Fu, J., 1990. Comprehensive Research on Permian in Guizhou–Guangxi Area and its Petroleum Potential. Science Press, Beijing (in Chinese). Shaw, H.F., Primmer, T.J., 1991. Diagenesis of mudrocks from the Kimmeridge Clay Formation of the Brae Area, UK North Sea. Mar. Pet. Geol. 8, 270–277. Shen, S., Wang, Y., Henderson, C.M., Cao, C.Q., Wang, W., 2007. Biostratigraphy and lithofacies of the Permian System in the Laibin–Heshan area of Guangxi, South China. Palaeoworld 16, 120–139. Shen, S., Zhang, H., Zhang, Y., Yuan, D., Chen, B., He, W., Mu, L., Lin, W., Wang, W., Chen, J., Wu, Q., Cao, C., Wang, Y., Wang, X., 2019. Permian integrative stratigraphy and timescale of China. Sci. China Earth Sci. 62, 154–188. Sholkovitz, E.R., Landing, W.M., Lewis, B.L., 1994. Ocean particle chemistry: The fractionation of rare earth elements between suspended particles and seawater. Geochem. Cosmochim. Acta 58, 1567–1579. Strasser, A., Hilgen, F.J., Heckel, P.H., 2006. Cyclostratigraphy – concepts, definitions, and applications. Newsl. Stratigr. 42, 75–114. Sun, Y., Lai, X., Jiang, H., Luo, G., Sun, S., Yan, C., Wignall, P.B., 2008. Guadalupian (Middle Permian) conodont faunas at Shangsi section, northeast Sichuan Province. J. China Univ. Geosci. 19, 451–460. Tostevin, R., Shields, G.A., Tarbuck, G.M., He, T., Clarkson, M.O., Wood, R.A., 2016. Effective use of cerium anomalies as a redox proxy in carbonate–dominated marine settings. Chem. Geol. 438, 146–162. Visscher, P.T., Reid, R.P., Bebout, B.M., Hoeft, S.E., Macintyre, I.G., Thompson, J.A., 1998. Formation of lithified micritic laminae in modern marine stromatolites (Bahamas): the role of sulfur cycling. Am. Mineral. 83, 1482–1493. Visscher, P.T., Stolz, J.F., 2005. Microbial mats as bioreactors: populations, processes, and products. Palaeogeogr. Palaeoclimatol. Palaeoecol. 219, 87–100. Wang, Y., Jin, Y., 2000. Permian palaeogeographic evolution of the Jiangnan Basin, South China. Palaeogeogr. Palaeoclimatol. Palaeoecol. 160, 35–44. Wang, L., Lu, Y., Zhao, S., Luo, J., 1994. Permian Lithofacies Palaeogeography and Mineralization in South China. Geological Publishing House, Beijing (in Chinese). Wang, X., Shi, X., 2008. Sedimentary characteristics and evolution of the Late Paleozoic Leye isolated carbonate platform in northwest Guangxi. J. Palaeogeogr. 10, 329–340 (in Chinese with English abstract). Waterhouse, H.K., 1999. Regular terrestrially derived palynofacies cycles in irregular marine sedimentary cycles, Lower Lias, Dorset, UK. J. Geol. Soc. 156, 1113–1124. Webb, G.E., Kamber, B.S., 2000. Rare earth elements in Holocene reefal microbialites: a new shallow seawater proxy. Geochem. Cosmochim. Acta 64, 1557–1565. Webb, G.E., Nothdurft, L.D., Kamber, B.S., Kloprogge, J.T., Zhao, J.X., 2009. Rare earth element geochemistry of scleractinian coral skeleton during meteoric diagenesis: a sequence through neomorphism of aragonite to calcite. Sedimentology 56, 1433–1463. Westphal, H., 2006. Limestone–marl alternations as environmental archives and the role of early diagenesis: a critical review. Int. J. Earth Sci. 95, 947–961. Westphal, H., Munnecke, A., 2003. Limestone–marl alternations: a warm–water phenomenon? Geology 31, 263–266. Westphal, H., Head, M.J., Munnecke, A., 2000. Differential diagenesis of rhythmic limestone alternations supported by palynological evidence. J. Sediment. Res. 70, 715–725. Westphal, H., Munnecke, A., Pross, J., Herrle, J.O., 2004. Multiproxy approach to understanding the origin of Cretaceous pelagic limestone–marl alternations (DSDP site 391, Blake–Bahama Basin). Sedimentology 51, 109–126. Westphal, H., Munnecke, A., Böhm, F., Bornholdt, S., 2008. Limestone–marl alternations in pelagic sea settings – Witnesses of environmental changes or diagenesis? In: Pratt, B.R., Holmden, C. (Eds.), Dynamics of Epeirics Seas. Geological Association of Cananda, pp. 389–406 Special Paper 48. Westphal, H., Munnecke, A., Brandano, M., 2008b. Effects of diagenesis on the astrochronological approach of defining stratigraphic boundaries in calcareous rhythmites: The Tortonian GSSP. Lethaia 41, 461–476. Westphal, H., Hilgen, F., Munnecke, A., 2010. An assessment of the suitability of individual rhythmic carbonate successions for astrochronological application. Earth Sci. Rev. 99, 19–30. Westphal, H., Lavi, J., Munnecke, A., 2015. Diagenesis makes the impossible come true: intersecting beds in calcareous turbidites. Facies 61, 1–7. Wheeley, J.R., Cherns, L., Wright, V.P., 2008. Provenance of microcrystalline carbonate cement in limestone–marl alternations (LMA): aragonite mud or molluscs? J. Geol. Soc. 165, 395–403. Wu, S., Yan, J., Liu, K., Yan, Y., 2016. Response of early Permian silisiclastic depositional

784

Marine and Petroleum Geology 111 (2020) 765–785

C. Su, et al. system to the advance of Gondwana glaciation in Southwestern Guizhou. Earth Sci. Front. 23, 299–311 (in Chinese with English abstract). Xiao, D., Tan, X., Xi, A., Liu, H., Shan, S., Xia, J., Cheng, Y., Lian, C., 2016. An inland facies–controlled eogenetic karst of the carbonate reservoir in the Middle Permian Maokou Formation, southern Sichuan Basin, SW China. Mar. Pet. Geol. 72, 218–233. Xue, W., Liu, X., Yan, J., Ma, Z., 2015. The origin of eyeball–shaped limestone from Maokou Formation (Mid–Permian) in Nanchuan region, Chongqing, Southwest China. Chinese J. Geol. 50, 1001–1013 (in Chinese with English abstract). Yan, J., Carlson, E.H., 2003. Nodular celestite in the Chihsia Formation (Middle Permian) of south China. Sedimentology 50, 265–278. Yan, J., Zhang, K., Carlson, E.H., 2000. Remarks on origin of sepiolite with evidences from replacement history of nodular celestite in the Chihsia Formation of South China. J. Mineral. Petrol. 20, 1–5 (in Chinese with English abstract). Yan, J., Munnecke, A., Steuber, T., Carlson, E.H., Xiao, Y., 2005. Marine sepiolite in Middle Permian carbonates of South China: Implications for secular variation of Phanerozoic seawater chemistry. J. Sediment. Res. 75, 328–338.

Yan, J., Ma, Z., Xie, X., Xue, W., Li, N., Liu, D., 2008. Subdivision of Permian fossil communities and habitat types in northeast Sichuan, South China. J. China Univ. Geosci. 19, 441–450. Yan, Y., Yan, J., Wu, S., 2015. Sedimentary records of Early Permian major glacial sea–level falls in southern Guizhou Province, China. Earth Sci. J. China Univ. Geosci. 40, 372–380 (in Chinese with English abstract). Yang, Z., Chen, S., Xu, J., 1988. Geological and prospective significance of the Early Permian, marine sepiolite–phase transformation in Pingxian–Leping Depression. Acta Sedimentol. Sin. 6, 91–99 (in Chinese with English abstract). Zhan, R., Jin, J., Liu, J., Corcoran, R., Luan, X., Wei, X., 2016. Meganodular limestone of the Pagoda Formation: A time–specific carbonate facies in the Upper Ordovician of South China. Palaeogeogr. Palaeoclimatol. Palaeoecol. 448, 349–362. Zhao, M.Y., Zheng, Y.F., 2017. A geochemical framework for retrieving the linked depositional and diagenetic histories of marine carbonates. Earth Planet. Sci. Lett. 460, 213–221.

785