Citation: | Rongcai ZHENG, Yanghui PAN, Can ZHAO, Lei WU, Renjin CHEN, Rui YANG. Carbon and oxygen isotope stratigraphy of the oxfordian carbonate rocks in Amu Darya basin. Journal of Earth Science, 2013, 24(1): 42-56. doi: 10.1007/s12583-013-0315-4 |
The carbon and oxygen stable isotope compositions of carbonate minerals are good tracers of sedimentary-diagenetic environments and formation of the basin. Hence, it is one of the most important topics in the isotope research of sedimentary rocks (Li and Wan, 1999). However, the carbon and oxygen stable isotope compositions of carbonate rocks are often superimposed results of deposition and diagenesis. Therefore, carbon and oxygen isotopes play an important role in geochemical tracers based on the application of deposition and diagenesis. For example, Chen and Chen (1994), Chen et al. (1995), and Zheng and Liu (1997) studied Devonian carbon isotope stratigraphic evolution and their relations to transgression-regression patterns in South China and Longmen Mountains. Using carbon, oxygen, and strontium isotopes, Zheng et al. (2008a, b; 2007) studied the genesis of dolomite reservoirs of Changxing Formation (Upper Permian) and Feixianguan Formation (Lower Triassic) in the northeastern Sichuan. Li et al. (2007) and Liu et al. (2007) studied the sequence stratigraphy of Triassic carbonate rocks in Southwest Guizhou and Upper Jurassic strata in Qiangtang Basin, respectively. Based on the detailed petrologic results (Zhang et al., 2010) and the application of carbon and oxygen stable isotopes, the carbon and oxygen isotope geochemistry characteristics of Oxfordian carbonate rocks were focused on in the Amu Darya Basin, Turkmenistan. It will provide important geochemical information for the analyses of the sea-level change, isotope stratigraphic correlation, and depositional-diagenetic environment in the Amu Darya Basin.
Amu Darya Basin is a large Mesozoic superimposed sedimentary basin in southeastern Tulan Platform. The basin is divided into three structural layers, including basement, intermediate layer, and platform cover (Meisel et al., 1995). The basement consists of Paleozoic igneous and metamorphic rocks at varying burial depths. Based on the structural morphology of the cover, the Amu Darya Basin comprises a number of large structural units, including Kopetat piedmont depression, central Karakym uplift, Charjou terrace, and so on (Fig. 1). Two groups of northwest- and northeast-striking faults controlled the structural framework and the distribution of sedimentary cover. The intermediate layer above the basement is composed of Permian-Triassic terrigenous clastic rocks. The Lower to Middle Jurassic coal-bearing clastic rocks of coastal plain and littoral-neritic sea had excellent source rocks and great resource potential. The widespread sedimentary cover is composed of Callovian-Oxfordian (Middle to Upper Jurassic) marine clastic rocks and carbonate rocks, extremely thick Cretaceous to Paleogene evaporites. In Callovian-Oxfordian, good carbonate reservoirs generally developed; the Cretaceous-Paleogene evaporites were high-quality regional cap rocks; and they together with Lower to Middle Jurassic strata formed a very favorable source-reservoir-cap configuration in time and space.
The Callovian-Oxfordian (Middle-Upper Jurassic) lithological association in the Amu Darya Basin is complex. It includes migmatites of shallow marine facies, shallow-abyssal carbonate rocks, and evaporites. The strata uncomformably overlapped the underlying Middle Jurassic coal-bearing clastic rocks of transitional facies and were continuous with the overlying Kimmeridgian and Tithonian (Upper Jurassic) extremely thick gypsum-salt rocks. The Callovian-Oxfordian strata are the most important gas producing horizons in the Samandepe Gas Field. They are composed of limestones and gypsum-salt rocks of basin, ramp, open platform, restricted platform and evaporative platform facies and can be subdivided into seven lithologic units from bottom to top: the under reef (XVa1), dense layer (Z), reef layer (XVa2), above reef (XVhp), massive limestone (XVm), bedded limestone (XVp), and limestone-gypsum layer (XVac) (Fig. 2). The strontium isotope of rudists from the massive limestone layer (XVm) is dated as 157.2 Ma (Zheng et al., 2011). Reservoirs mainly developed above the Oxfordian reef (XVhp). On the top, the limestonegypsum layer was the dense cap rock.
Fifty-three samples were selected for carbon and oxygen isotopes, as well as Fe, Mn, and Sr trace elements in our work. The sample types and analysis results are shown in Table 1. The samples were obtained from the cores of borehole Sam53-1 in the Samandepe Gas Field, with sampling depth in the range of 2 366-2 658 m and mainly covering the Oxfordian strata (Fig. 2). The C and O isotopes were performed in Geological Laboratory, Research Institute of Exploration and Development, Southwest Oil-Gas Branch Company, CNPC. The analytical instrument is MAT-252 EM mass spectrometer made by Finnigan Company, with ±0.2‰ analytical precision and PDB standard. The Fe, Sr, and Mn trace elements were performed in the Analysis Center, Chengdu Institute of Multipurpose Utilization of Mineral Resources, Chinese Academy of Geological Science. The test instrument is ICP-AES 5300V made by Thermo Electron Corporation in USA. The test standard is Y/T05-1996 < ICP broadspectrum>, with 0.001% detection limit and ±0.000 5% error. The accuracy of the test instrument can meet research requirements.
From statistical data in Table 1, the following characteristics are obtained for the carbon and oxygen isotopic composition of different types of Oxford carbonate samples: ① The δ13C (4.08‰) and δ18O (-1.913‰) of normal marine limestones represented by micrite limestones and grain micrite limestones are almost the same to the δ13C (3‰-4.5‰) and δ18O (-1.17‰ - -3.0‰) of Oxfordian normal marine limestones of the ancient Mediterranean reported by Bartolini et al. (1999); ② the δ13C (4.039‰-4.433‰) and δ18O (-1.313‰ - -1.904‰) of most reef limestones and granular limestones are close to contemporaneous normal marine limestones; ③ the δ13C (4.038‰-4.158‰) of dolomitic limestones, crystalline dolomites and hydrothermal calcites are also consistent with contemporaneous normal marine limestones, but the δ18O (-3.18‰ - -4.072‰) are more negative than normal marine limestones; ④ 2 chalk samples have the lowest δ13C (2.99‰) and δ18O (-4.86‰).
The above-mentioned results show that normal marine limestones represented by micrite limestones and grain micrite limestones formed in the normal marine environment, but several other genetic types are more or less affected by external fluids in diagenesis, typically characterized by negative δ18O excursion of some samples and gradually increased negative δ18O excursion from micrite limestones to granular limestones, reef limestones, calcitic dolomites, crystalline dolomites, and hydrothermal calcites with increasing diagenetic intensity. Consequently, when analytical data in Table 1 are used for geological interpretation, the applicability of the sample, especially the representative reliability of the original seawater of depositional environment, must be assessed. In many existing research results, Mn and Sr content and Mn/Sr ratios of the samples are generally important indicators of its representative of seawater. For example, Korte et al. (2003) considered the samples with less than 100×10-6 Mn content were good in the study of isotope stratigraphy of the Alpine Triassic marine limestones. Kaufnan et al. (1992) believed that the isotopic composition with only Mn/Sr < 2 and δ18O>-5.0‰ would be of value when using strontium and carbon isotope changes to inverse sea level and ancient climate change. The samples used for paleoenvironmental and isotope stratigraphic analysis must meet the following conditions: ① to avoid the part of obvious diagenetic alteration as much as possible and no strong recrystallization in thin sections; ② high Sr (≥200×10-6) and low Mn content (< 100×10-6); ③ Mn/Sr ratios < 1; and ④ δ13C in the range of 3‰-4.5‰ and δ18O≥-3.0‰ deviation (Bartolini et al., 1999). The applicability and reliability of all the samples were assessed in accordance with the above conditions, and the results can be divided into three cases: ① 24 samples can meet the above four conditions (Table 1, sample marked with a), δ13C is 3.78‰- 4.83‰, with an average of 4.186‰, δ18O is -0.11‰ - -2.91‰, with an average of -1.523‰, within the carbon and oxygen isotopic composition of Oxfordian normal marine limestones in the ancient Mediterranean by Bartolini et al. (1999). Therefore, they can reliably represent the carbon and oxygen isotopic composition of the original seawater and can be used for isotope stratigraphic and paleoenvironmental analysis; ② 15 samples can meet most conditions (Table 1, samples marked with ab) but had positive δ13C excursion to varying degrees, negative δ18O excursion, and slight Sr loss, thus they had worse representation of carbon and oxygen isotopic composition of the original seawater. Five samples among them relatively enrich in 13C but meet the requirements of 18 O and Sr contents so still have a significance of good isotope stratigraphic and paleoenvironmental analysis; ③ the remaining 14 samples cannot meet the above conditions (Table 1, samples marked with b), and they only can be used for diagenetic analysis.
In a balanced diagenetic system, the migration of δ13C needs pore solution at least 1 500 times of their own volume, and the same change of δ18O needs a pore solution only five times more than their own volume (Li and Liu, 1996). Therefore, carbon isotopes of carbonate rocks reflect more effective information of the ancient sea-level changes than oxygen isotopes (Veizer et al., 1986). The cementation, fresh water leaching, and replacement, recrystallization, as well as filling in burial diagenetic process are the main destruction of the original carbon isotope preservation. The published literatures have proved that the oxidation, burial flux, and burial rate of organic carbon are the most important factors affecting the carbon isotope of carbonate rocks (Yin and Ni, 2009; Prokoph et al., 2008; Wang and Bai, 1999; Kaufman et al., 1997; Kaufman and Knoll, 1995; Derry et al., 1994, 1989). The burial flux and burial rate of marine organic carbon are significantly controlled by the sea-level change. For example, the continental area become smaller and the ocean area expand in the period of sea-level rise, the burial flux of bio-organic carbon increases with the rising sea levels and the flux of marine organic carbon significantly reduces, CO2 dissolved in seawater contains more 13C, then the carbonates balanced with the seawater are also rich in 13C, and the corresponding carbonate sediments have relatively big δ13C values (Niebuhr and Joachimski, 2002). In contrast, the continental area increase and the ocean area reduce in the period of continuing sea-level decline, large amounts of organic matter on land are exposed and oxidized to 12C-rich CO2 and brought into seawater because of erosion; meanwhile, marine organisms reduce, the burial rate of organic matter reduces, and oxidation enhances, resulting in large amounts of 12C-rich CO2, which is released to the seawater and make δ13C in sediment decrease.
According to the plots of carbon isotope stratigraphic data of 29 samples, which can be used for isotope stratigraphy and paleoenvironmental analysis (Fig. 3), the cyclicity of Oxfordian carbon-isotope stratigraphic curve is consistent with Oxfordian paleoenvironmental evolution in Amu Darya Basin and can be compared with the Oxfordian carbon-isotope stratigraphic curve in the ancient Mediterranean by Bartolini et al. (1999), the change trend of the global sea-level by Prokoph et al. (2008), and the sea-level of the South China (Zuo, 2003). The similarity of the evolution trend of carbon-isotope stratigraphy proved that the samples well saved the isotopic composition of the original seawater and can provide the basis for the Oxfordian sea-level changes, paleoenvironmental analysis, and isotope stratigraphy correlation in Amu Darya Basin.
The Oxfordian sea-level change history determined by the carbon isotope stratigraphic curve in the Amu Darya Basin can be divided into three evolutionary stages (A, B, and C) (Fig. 3): ① A stage, equivalent to Early Oxfordian, developed in the above reef (XVhp) to the lower massive limestone (XVm). The depositional environment changed from the foreslope transitionally to the reef on the platform margin. The relative sea-level overall declined, forming shallowing-upward sequence. The δ13C decreased from 4.28‰ to 3.78‰; ② B stage, equivalent to MidOxford, developed in the middle massive limestone (XVm) to the upper bedded limestone (XVp). The relative sea level rose. The sedimentary environment was alternating reefs and shoals on the platform margin, hence, developed the most important reef-bank reservoirs in Amu Darya Basin. The sea-level fluctuations were more frequent to late B stage, and the sedimentary environment was alternating shoals of open platform and subtidal hydrostatic mud, dominated by thinner bioclastic beach, sandy beach, and oolitic beach, locally with patch reef reservoirs. In this stage, carbon isotopes showed continuing positive excursion, δ13C increased from 3.78‰ to 5.98‰ with more than 2.2‰ amplitude, indicating that controlling factors was the rising sea level. Because of stable sedimentary environment, burial rate of organic carbon increased with rising sea levels, and organic carbon in seawater decreased, CO2 dissolved in seawater contained more 13C, and carbonate rocks balanced with the seawater were also rich in 13C. Combined with the geological background, it could be interpreted as the sedimentary geochemistry response to the positive carbon-isotope excursion in the Mid-Oxfordian continuing transgression. Apparently, the extensive development of the Oxfordian reefs and banks in Amu Darya Basin is closely related to the transgression; ③ C stage, equivalent to Late Oxfordian, developed in the upper massive limestone (XVm) to limestonegypsum layer (XVac). The depositional environment changed from the open platform through restricted platform into evaporative platform and then transitioned upward into the more extensive Sabkha environment. The lithology was interbedded dense micrite limestones and grain micrite limestones; some limestones had anhydritigation and penecontemporaneous dolomitization, localized with thin gypsum-salt rocks, reflecting that the relative sea level declined, and the water circulation was limited and salinized; the δ13C values decreased from 5.98‰ to 4.64‰ but still larger than the average of the normal marine limestone.
From carbon isotope evolution trend in Fig. 4, it can be seen that there are two obvious global positive carbon-isotope excursions in Early and Middle Oxford (Late Jurassic), one is Mid-Oxfordian cordatum zone, and the other is from Mid-Oxfordian plicatilis zone to transversarium zone; δ13C was the biggest at early transversarium. Oxfordian carbon isotope stratigraphic curves and Mid-Oxfordian positive excursion event in Amu Darya Basin can be compared to SE France and Paris Basin in the northern France (Lavastre, 2002), Swiss Jura Mountains (Gygi, 1999), and western Carpathians in Poland (Wierzbowski, 2002) (Fig. 4). Previous studies believed that the Mid-Oxfordian positive carbon-isotope excursion event was closely related to sustained global transgression and biological prosperity, increased organic carbon burial flux, and intermittent decline of atmospheric carbon dioxide partial pressure. Meanwhile, global warming led to the widespread deposition dominated by carbonate rocks, for example, Louis-Schmid et al. (2007) studied the Oxfordian carbon isotope stratigraphy in France and Swiss and believed that the large-scale development of carbonate platform buildups and reefs in this period are closely related to sea-level rise. Apparently, the MidOxfordian sustained extensive transgression laid the foundation for the prosperity of reef-building organisms, represented by rudists and corals and a wide development of reef-bank reservoirs in Amu Darya Basin, which also provides another example for studying Oxfordian (Late Jurassic) global transgression event and the resulting large development of reefs and banks and high-speed burial of organic carbon.
As mentioned earlier, δ18O has a greater sensitivity to negative excursion in diagenesis than δ13C. Because the negative δ18O excursion is usually explained by fresh water mixing and temperature effects, people have different point of views about the question on which factor plays a dominant role in the negative δ18O excursion of original sediment while recovering the diagenetic environment. For example, when people explained the dolomite genesis of Feixianguan Formation in northeastern Sichuan Basin, different scholars considered that the dolomites were replaced by brine seepage refluction (Luo et al., 2006), replaced by mixed water (Wang et al., 2009), or replaced by hydrothermal fluid (Zheng et al., 2008a, b) for the same negative δ18O excursion. Based on the δ18O values of Late Jurassic seawater and fossils provided by Bartolini et al. (1999), the present study dealt with the degree of diagenetic alteration and diagenetic fluid properties of carbonate samples, and the following understandings are obtained.
① The average δ13C of different types of samples (2.99‰-4.433‰) have little change and positive excursion (Fig. 5 and Table 1), whereas the δ18O (-1.904‰- -4.86‰) vary in a large range and show more negative excursion from micrite limestones to granular limestones, reef limestones, dolomitic limestones, crystalline dolomites, hydrothermal calcites, and chalks, indicating that the fluid restricted in the formation was gradually consumed as the diagenesis enhanced, while the sediments transformed.
② The δ18O values of various types of samples range mostly near the δ18O values of Late Jurassic seawater in the ancient Mediterranean by Bartolini et al. (1999), micrite limestones, grain micrite limestones, and granular limestones had the largest distribution among the rock types (Fig. 6). According to isotope fractionation principles, the isotopes of seawater with evaporation effect have more positive excursion than normal seawater, so δ18O values of the fluid originated from seawater or evaporating seawater are close to or slightly greater than contemporaneous seawater. For example, most samples located in area C in Fig. 6 are micrite limestones and grain micrite limestones in the restricted platform, of which relatively enriched 18O are related to the higher seawater salinity of depositional environment than the open sea. Most samples located in area B in Fig. 6 are micrite limestones, grain micrite limestones, granular limestones, and reef limestones of the open platform and platform margin. Most samples located in area A are calcitic dolomites, crystalline dolomites, and hydrothermal calcites precipitating from the formation water and chalk suffering from intense diagenesis. Overall, samples in area B, C suffered from weaker diagenetic alteration, and it can be evidenced by the fact that micrite limestones and grain micrite limestones are dense, with no significant pores and fractures and weak recrystallization (Fig. 7a). Granular limestones and reef limestones that have also suffered from relatively weaker diagenesis are still in the middle diagenesis dominated by cementation, slight compaction, and dissolution, adapting to the characteristics of well-preserved original intergranular pores, skeletal, and framework pores (Figs. 7b-7d), which also saved the partial information of the original seawater to a large extent. The original structures of samples in area A were completely destroyed in the diagenetic process (Figs. 7e, 7f).
③ Because the buried conditions had higher temperature than surface conditions, thermal fractionation made stable isotopes into the mineral lattice, which led diagenetic fluids under the burial to have more negative δ18O values than the contemporaneous seawater. Therefore, compared with the δ18O values of contemporaneous seawater, dolomitic limestones, crystalline dolomites, hydrothermal calcites, and chalk samples, which have significantly more negative δ18O excursion than the range of normal seawater (Fig. 6), are believed to be the products of the burial diagenesis. Besides, temperature effect was the dominant factor in this process.
④ In equilibrium conditions, the oxygen isotopic composition of authigenic carbonate minerals precipitating from seawater and lake water is a function of temperature and oxygen isotope composition of the water. Gasse et al. (1987) gave the following formula on the basis of previous studies
where SST (℃) is water temperature at the precipitation of carbonate minerals, δ18Oc is the δ18O value (PDB) of the measured samples, and δ18Ow was the oxygen isotopic composition of water. The formula is mainly used to reflect the formation temperature of carbonate minerals, but δ18O values of minerals will change in the geological history, making the application of the formula restricted; but this calculation still has a reference value to determine the extent of diagenetic alteration (Liu et al., 2006). According to Bartolini et al. (1999), the δ18O of the Oxfordian Mediterranean seawater represented by normal marine limestones was between -1.17‰ and -3.0‰, with an average of -2.085‰. Calculated water temperature at the formation of the dolomitic limestones, crystalline dolomites, hydrothermal calcites, and chalk is between 15.41-20.97 ℃ (Table 1), which is significantly higher than the calculated 15.16-20.62 ℃ of micrite limestones and 14.94-17.25 ℃ of granular limestones and reef limestones. From another point of view, temperature was the dominant factor in the negative δ18O excursion effects in the diagenetic environment.
⑤ Combined with carbon isotope stratigraphy and compared to δ13C and δ18O values of Late Jurassic seawater in the ancient Mediterranean, the average 13C values of all kinds of samples (2.99‰-4.87‰) are positively excursed than the δ13C values of contemporaneous normal seawater in the ancient Mediterranean (3.0‰-4.5‰), indicating that the formation water in the diagenetic process is relatively saltier than the contemporaneous normal seawater, and 13C values were in an enriched state. It has been proven that the role of fresh water, oxidation, or fermentation of organic matter can lead to depleted 13C (Luo et al., 2006; Veizer et al., 1999; Wang and Bai, 1999), and if hydrocarbons are present in the samples, 13C values appear to be more substantial than negative excursion, so from the positive δ13C excursion and negative δ18O excursion of all kinds of samples, it is easy to determine that the diagenetic environment is far from freshwater, and dolomitization and dissolution were completed before large numbers of hydrocarbons charge. This feature further proved that dolomitization and precipitation of hydrothermal calcites occur red in the closed diagenetic environment where the temperature is the controlling factor.
⑥ δ13c and δ18o of hydrothermal calcites are very close to dolomitic limestones and crystalline dolomites. combined with homogenization temperatures of gas-liquid two-phase fluid inclusions and salinity measurement results of four hydrothermal calcite samples filling solution vugs and fractures (table 2), it can also be determined that dolomitization and precipitation of hydrothermal calcites occurred in the diagenetic fluids with the nature of hot brine in a certain burial depth, temperature, and pressure conditions, were the products of varying water-rock interaction between the same diagenetic fluids with different objects in different stages in hot brine system.
⑦ It is worth mentioning that two chalk samples have the lowest δ13C and δ18O values. Relative to normal marine limestones, their δ18O have more negative excursion than δ13C (Figs. 5 and 6). The two chalk samples were taken from the gas-bearing water layer, so it is believed that the chalkification of limestone reservoirs was related to oilfield water with poor 13 C and 18O, and δ18O values have more significant thermal alteration effects of negative excursion in the chalkification process.
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