Aptian to Campanian sediments from the Western offshore to Central Orange Basin were studied by integrating molecular geochemistry, inorganic and isotopic studies to recognize their geochemical characteristics via the reconstruction of the Orange basin’s paleoweathering, paleosalinity, paleovegetation, paleoclimate, and tectonic records. Molecular analyses of both aliphatic and aromatic compounds reveal an input dominantly from a marine source. The source rocks accumulated in a reduced, anoxic, saline water column. Based on various biomarker proxies and vitrinite reflectance data, some samples are thermally mature to produce petroleum, while others are not. According to the V/Ni ratio, samples from the Orange Basin in South Africa are mainly anoxic, with only a few samples ranging from suboxic to anoxic. This is congruent with biomarker and isotope analyses that further indicate the presence of marine-derived source rocks with some terrestrial remains generating hydrocarbons. The investigated sediments are made up of intermediate igneous rocks that have undergone moderate chemical weathering. Geochemical figures on tectonic setting discriminant function diagrams revealed a continental rift of passive margin settings. As a result, the extrapolated crustal processes are directly analogous to the genesis and evolution of the Orange Basin, demonstrating Gondwana’s breaking up and the opening of the Atlantic Ocean Margin.
During the last thirty years, the South African government has been motivated to explore the Orange Basin due to the discovery of valuable petroleum on the southern Atlantic margins (Jungslager 1999; Hartwig et al. 2012; Adekola et al. 2012). The huge Ibhubesi gas reserve off South Africa as well as the Kudu gas field in South Namibia are the result of ongoing hydrocarbon prospecting (PASA 2003). There have been some reports on the geochemical attributes of source rocks from the Orange Basin (Jungslager 1999; Van der Spuy 2003). According to Jungslager (1999) and Van der Spuy (2003), the Orange Basin has primarily three source rock intervals.
Organic geochemistry is an important tool for identifying source rocks and classifying crude oils into families both at the beginning and later stages in oil exploration (Peters et al. 2005; Akinlua et al. 2007a, 2007b; Adekola et al. 2012). A thorough geochemical analysis of source rocks can reveal information about the hydrocarbons they produce (Adekola et al. 2012). The composition and long-term preservation of organic matter deposited in the marine domain throughout the Orange Basin is related to deposition system stacking and, as a result, the depositional systems tracts in which it was deposited (Adekola et al. 2012). Most geochemical analyses rely heavily on data obtained from bitumen extraction. This yields biomarkers from both aliphatic and aromatic fractions that could be utilized to deduce organic source input, paleoredox condition, maturity, and age determination as well as relating their provenance (Peters et al. 2005). Therefore, the quantity, quality and thermal maturation of source rocks are pivotal in petroleum exploration. In order to answer some questions such as organic matter origin, depositional condition, thermal maturation, type and quality of source rocks, researchers ought to work on many representative samples extending up to the deeper Orange Basin. In characterizing the source rocks, these conventional methods may be related to unconventional approaches such as nuclear magnetic resonance (NMR), Fourier transform infrared spectroscopy (FTIR), and thermogravimetric analysis (TGA).
In petroleum exploration, trace elements are also used to delineate the emergence, depositional environment, thermal maturity of organic matter, as well as pathways of oil movement (Lewan 1984; Akinlua et al. 2015, 2018). Metal concentration ratios (V/(V + Ni)) is used to identify organic matters within sediments, such as siliceous versus phosphatic shales (Akinlua et al. 2010). The types of organic matter and depositional environment, according to Lewan (1984), have a significant impact on the amount of trace elements in source rocks. As a result, the trace element content of source rock has become a valuable geochemical tool for evaluating source rock. The concentration of trace elements certifies the source rock type (Barwise 1990). A reliable paleoenvironmental marker, the trace element is not inhibited by maturation and is used to depict depositional age in source sediments as well as likely time to generate hydrocarbons. Cumming et al. (2014) determined the depositional age of source rocks and the timing of petroleum generation in the Uinta Basin using rhenium–osmium geochronology (Re-Os). Some research has been conducted on crude oil and condensates, as well as on the source rock geochemistry of the South African basin (Davies 1997; Jungslager 1999; Vander Spuy et al. 2003; Akinlua et al. 2010) but scarce inorganic geochemical work undertaken on whole rock samples. The study of stable carbon isotopes is used to investigate environmental issues that impact organic matter, like paleoclimatic evolution (Hentschel et al. 2016); provenances (Thornton and McManus 1994).
The integration of organic and inorganic geochemistry provides a multiscale perspective that allows researchers to answer some questions such as origin of organic matter, genetic linkage within the source rocks, deposition, thermal maturity, sediment age, type and quality of source rocks, mineral content, and chemical composition of the vast, unexplored Orange Basin post-rift sequence. In this context, molecular geochemistry, inorganic methods and isotopic studies were integrated to recognize their geochemical characteristics through the reconstruction of the Orange basin’s paleovegetation and paleoweathering. This approach could also be adapted to understand certain tectonic elements and prospectivity in other basins in the world. In addition, the characterization of some representative rock samples for elementary analysis resulted in a useful environmental and health implications database.
The geologic setting
The Orange Basin is an offshore passive continental boundary of Southwest Africa’s Atlantic Ocean (Fig. 1; Akinlua et al. 2011; Jungslager 1999) that is bounded to the north by the Kudu Arc and to the south by the Agulhas-Columbine Arc (Fig. 2; Hirsch et al. 2007; Muntingh and Brown 1993). The basin formed in the late Jurassic-Tertiary period (Brown 1995; Gerrard and Smith 1982; Muntingh 1993). It is located between latitudes 31–33.5 S and longitudes 16–17 E, stretching nearly 160,000 km2 extending from the continental shelf toward the deep offshore marine, as shown in Fig. 1 (PASA 2003; Broad et al. 2006).
Orange Basin seems to be larger than the other six (6) basins in South Africa put together (McMillan 2003). Its main source is the Orange River, which empties into the South Atlantic Ocean (Akinlua et al. 2011). Despite the Orange Basin being South Africa’s largest offshore basin, less than 40 wells have been drilled, equating to one well per 4000 Km2 (Muntingh 1993). An estimated 43,000 km2 of deep-water license area extends from Namibia’s offshore southwest volcanic passive margin (Paton et al. 2007) spanning from the upper Jurassic to the Holocene (Kuhlmann et al. 2010). Figure 2 illustrates the basin’s many onshore provinces, mobile belts, craton, and the Kalahari shield.
Source rock intervals in the Orange Basin are best explained by tectonostratigraphic events. The tectonic stages are as follows: late synrift, early post-rift, and drift succession (Jungslager 1999). These features resemble sedimentary trenches (Akinlua et al. 2010) made of source rock (Van der Spuy et al. 2003).
According to petroleum systems, they are also classified based on individual organic matter content, hydrocarbon type, and spatial distribution. Within the Orange Basin, three (3) most qualitative source rock intervals were discovered. These major source rock intervals are as follows:
Hauterivian synrift lacustrine source rocks (Type I)
Pelagic transition sediments from the Barremian to the Lower Aptian (Type II/III) (Jungslager 1999).
Marine source rocks from the Aptian drift stage (112-103 Ma)
Cenomanian–Turonian Type II drift marine source rocks (Vander Spuy et al. 2003).
The Hauterivian synrift source rock is only found in the Orange Basin at the AJ1 well’s half-graben (Barton et al. 1993). The source rock aids in the correlation and reconstruction of the environment up to the South American margins (Paton et al. 2007). Age-wise, Barremian rocks are the second category of source rocks, which is found between the transition–drift tectonic stage in the Aptian anoxic environment (Van der Spuy et al. 2003). It penetrates the AK-half-graben, as evidenced by the AK wells that produce gas. Likewise, the Barremian to Lower Aptian source manifests at OA1 and is oil prone (Jungslager 1999). The source rocks are likely to enrich Namibia’s Kudu gas and the Ibhubesi fields (Kuhlmann et al. 2010).
Finally, within the AK1’s half-graben, the Turonian drift phase source interval appears, presumably producing gas and oil (Van der Spuy et al. 2003). Fluvial and deltaic sandstones with conglomerate from the Paleozoic Karoo section and underlying basement rocks serve as reservoir rocks in the rifting region. Fluvial sandstones and floodplain sediments are found in the drift stage (Jungslager 1999). The Orange Basin contains a variety of play types, including rift plays, which are lacustrine sandstones that trap hydrocarbons from organic-rich claystone. Another significant play is the synrift sediments with drift plays, which include early Cretaceous aeolian sandstone plays, Albian incised valleys, shallow to deep-water structural sediments, and rolling over anticlines (Vander Spuy et al. 2003).
Sampling and analytic techniques
Twenty shales were collected from cutting samples of the Orange Basin in four-wells KD-1, KE-1, KF-1, and KH-1 (Fig. 1). The initial geochemical approach began with cleaning, removing oxidized surfaces, followed by crushing and grinding the shale to a fine powder with a pestle and mortar. About 5–10 g of the powdered sample was extracted for 72 h by using Soxhlet apparatus to evaluate extractable organic matter (OM) concentrates as well as the hydrocarbon content of the examined sediments. In a glass jar, 93 percent dichloromethane was mixed with 7 percent methanol before being poured into an empty solvent bottle and thoroughly shaken. A 250-ml conical flask and Soxhlet extractor apparatus were rinsed with solvent to neutralize any contaminant and then oven-dried. About 150 ml of the solvent was poured into the conical flask, and then alumina was added to avoid bumping and immediate evaporation of the solvent. Copper metal was also added into the conical flask of the solvent to absorb sulfur contained in the shale sample. The conical flask containing the solvent, anti-bump alumina, and Cu was mounted on the Soxhlet extractor, which housed the sample in the thimble and was connected to a condenser for 72 h of mantle heating. The extracted organic matter was separated using liquid column chromatography into saturated hydrocarbons, aromatics, and hetero-NSO fractions, which were then dried at room temperature before being investigated using gas chromatography (GC) or gas chromatography mass spectrometry (GC/MS).
Molecular approach and isotope method
The GC and GCMS were used on the aliphatic and aromatics to show the distribution of n-alkane, isoprenoid (Fig. 3a), hopane (Fig. 4a), and sterane (Fig. 5a) as well as some aromatic hydrocarbon fractions (Fig. 6a, b). Saturated fractions were mixed with hexane and analyzed using the gas chromatographs of the Agilent 6890 N series. A gaseous spectrometer FID coupled with an HP-5MS capillary was used for gas chromatography analysis, with intervals of 40 to 300 °C relative to 4 °C per minute, toward 300 °C for 30 min. Similarly, an Agilent V5975B-MSD mass spectrometer outfitted with gas chromatographs and ion sources was adjusted for gas chromatographic mass spectroscopic analysis. By scaling the peak heights of the n-Alkanes, isoprenoids and aromatic hydrocarbon parameters, Table 1; terpane, Table 2; and sterane/diasterane, Table 3 mass fingerprints, biomarker imprints were adapted for molecular geochemical identification. Individual compounds were derived based on retention time and compared to previous work’s mass spectrum (Adebanji et al. 2014). The GCMS was performed at the Geology Department, University of Malaya. Peak heights from ion-chromatograms were integrated to calculate aromatic hydrocarbon ratios (Van Aarssen et al. 1999). Agilent ChemStation software was used to process the data collected. For m/z 85, 191, 217 (branched/cyclic) and m/z 178, 184, 192 aromatics, the mass spectrometer was chosen. The bulk stable carbon isotope (13C/12C) of the whole rock sample was determined using an isotopic mass spectrometry apparatus and the combustion method (Sofer 1980). According to the global standard, the outcome was a delta, represented by the Pee Dee Belemnite (PDB).
Inorganic geochemical methods
According to Pi et al. (2013), the whole rock samples were pulverized and digested to examine some major and trace elements. On a finely crushed sample, an X-ray fluorescing spectrometer (XRF) investigation was carried out using a P-analytical Axios mAX 4 kW sequenced XRF spectrometer. An indestructible wavelength dispersive X-ray fluorescing spectrometer was used to determine major and trace oxides (WDXRF). 0.50 g of samples were placed in a Teflon beaker and dehydrated overnight at 105 °C to determine the concentration of minute traces and rare earth elements. The shales were dampened with milliliters of de-ionized aqueous solution; 5 ml HNO3 was added and placed on a hotplate at 150 °C just before drying; 10 ml HF was added, followed by 4 ml HClO4. The solution was heated to 200 °C using a crystalline mush until it was nearly dry. The settled and cooled samples were digested in a fume hood with 10 ml of 5 M HNO3.The solutions were diluted using de-ionized water in 50-ml volumetric flasks, and ultra-pure water was used to dilute the digested samples up to 100 times. 10 ppm analytic-concentration-standard element solutions were used for calibration, and the detection limit was set at less than one ppb. Trace elements were examined using the Agilent Technologies 7500S-ICP-MS.
Result and discussion
Molecular properties and organic matter as a source of input
Normal alkane containing pristane and phytanes
The fragmentogram of some of the samples studied, such as KD1-2250 and KH1-4232, contain low-to-moderate humps. Though, the n-alkanes extend further than C31, the humpy characteristics of some samples indicate low biodegradation (Volkman 1986). The fragmentograms show unimodal distributions of n-alkane and isoprenoids, additionally, low-to-intermediate-weight compounds predominate (n-C16-n-C26) (Fig. 3a), with high Ph peak illustrating reducing peaks as the number of n-alkane carbon atoms increases (Fig. 3a). In view of these characteristics, a marine carbonate source is more likely to contain algal-derived organic matter (Fig. 3b, d). Based on this data combined with Pr/n-C17 against Pr/Ph cross plot (Fig. 3b), oils are expected to generate and expel from marine carbonate source rock (Fig. 3d). Low-to-moderate-weight compounds with remarkable + n-C26 carbon atoms yield waxiness indexes ranging from 0.44 to 8.46 and values of the CPI range from 0.73 to 1.15 (Table 1). This evaluation is typical of marine shales with mixed algae that have terrestrial organic matter inputs (Fig. 3c, e). Comparison of Pr/n-C17 vs Ph/n-C18 in cross plots (Fig. 3c) show that algae-reducing environments predominate. These source rocks were genetically obtained via the same carbonate-rich lithofacies accumulated under a reducing to an oxidizing setting signifying primitive deposition of marine with mixed organic remains under reducing anoxic settings of the source rocks (Fig. 3b, d). Pr/Ph ratios range from 0.10 to 0.58, indicating a high aquatic contribution (Lijmbach 1975) accumulating under anoxia (Makeen et al. 2015). The ratio of Pristane to Phytanes is commonly used to depict paleoxygenation settings (Escober et al. 2011), which is higher with greater thermal maturation (Connan and Cassou 1980) and more terrestrial remains. Similarly, as salinity increases, Pr/Ph ratio falls (Peters et al. 2005). Lesser Pr/Ph (0.10–0.58), higher Ph/n-C18 (2.56–20), and moderate C.P.I. (0.64–1.15) under anoxic, saline water column with some terrestrial input is the source of rock on the western coast, according to the extracts (Fig. 3d, e; Peters et al. 2005; Bechtel et al. 2012). Comparison of Pr/n-C17 vs Ph/n-C18 in cross-plots (Fig. 3c) shows that algae-reducing environments predominate. These source rocks were derived genetically from the same carbonate-rich lithofacies that accumulated from a reducing to an oxidizing environment. Under the reducing anoxic settings of the source rocks, primitive deposition of marine with mixed organic remains occured. Based on this data combined with Pr/n-C17 against Pr/Ph cross-plot (Fig. 3b), hydrocarbons are expected to generate and expel from marine carbonate source rock.
Tricyclic terpanes found in source rock samples were used to characterize the rocks, correlate oils and extracts, and determine thermal maturation (Zumberge 1987; Peters et al. 2005). The extent of anoxic in mixed sediments is defined by C22T vs C21T, C24T vs C23T, C25T vs C26T, and C24Tet/C26T ratios (Table 2; Zumberge et al. 2005; El Diasty et al. 2016). Similarly, increasing marine source input are featured by cross-plots, as shown in Fig. 4b, c. Hopanoids (Fig. 4a) were made up of C29 norhopane, C30 hopane, some C31–C35 homohopanes, G/C30, and Tm. In all samples, C30 hopane predominates over C29 norhopane, as shown in Fig. 4a and Table 2, except for KH1-1565, which has a C29 norhopane/C30 hopane ratio of 1.01. This single sample was taken from an organic-rich carbonate evaporite rock (Zumberge 1984; Connan et al. 1986; Peters et al. 2005). C30H has a higher proportion than C29NH (C29-norhopane/C30H ratio of less than 1) characterizing abundant clay materials in the rocks (Gurgey 1999), as confirmed by C29 diasterane/R29 with 87 percent of analyzed samples greater than 0.2, indicating a suboxic deposition condition (Makeen et al. 2015; Moldowan et al. 1994; Peters et al. 2005). Thus, extracts have moderate-to-high amounts of pentacyclic and tricyclic terpane with less tetracyclic as determined by C24Tet/C26T ranging between 0.67 and 2.27; C23T/C24T ranging between 1.24 and 2.50, indicating reasonable terrestrial matter contributions (Fig. 4b). The high preponderance of C23T over C24T (C23T/C24T) 1.15–2.50 confirms microbial action within the examined samples, indicating a dysoxic environment. Furthermore, tricyclic terpane ratios based on the number of carbon atoms have been identified as a means of better understanding source input (Peters et al. 2005). The existence of average to high C24T/C23T values ranging from 0.40 to 0.87 indicates the presence of both marine and terrestrial matter (Marynowski et al. 2000). C21/C23T less than 0.5 indicates that the rock came from the sea (Quihua et al. 2011). The C21T/C23T ratio in the examined samples ranges from 0.10 to 0.50, with only one sample having 1.86 (KD1-2245). The source rocks were formed in a marine anoxic environment with terrestrial input (Fig. 4b). The ratio of C26T to C25T was modified to distinguish between bacterial and terrestrial plant matter (Volk et al. 2005). Low-to-medium C26T/C25T values range from 0.57 to 1.67 in this shaly sediment, with 87 percent greater than 0.9 signifying a mixture of algal and bacterial organic materials making up slightly terrestrial input (Volk et al. 2005). A high C26T versus C25T ratio, a low–medium C24T/C23T ratio, and a low-to-medium C22T/C21T ratio indicate lacustrine mix primal matter (Aquino et al. 1989; Kruge et al. 1990; Peters et al. 2005; Volk et al. 2005; Makeen et al. 2015). Some carbon isotope results support this as well (Fig. 13). The C31-22R/C30H ratio is primarily high, exceeding 0.25 in the marine environment and falling below 0.25 in the lacustrine environment of deposition (Peters et al. 2005). C31-22R/C30-H for investigated shales is between 0.21 and 0.46, with carbonate and shaly marine source input accounting for 80% of examined samples and lacustrine influence accounting for 20% (Fig. 4c; Peters et al. 2005).
Furthermore, aromatic fractions show low levels of aromatic compounds in some samples, indicating that the source rock depositional environment is marine (Radke et al. 2000). Lower DBT/P ratios indicate the incorporation of reducing primal sulfate matter (Hughes et al. 1995). Less DBT vs. P (Table 1) suggests algae with a minor terrestrial organic input. Gammacerane is present in all the shales studied, though in trace amounts in some (Table 2). It is made by reducing tetrahymanol (Venkatesan 1989). Tetrahymanol is produced by bacterivorous ciliates that live in the intervening oxic and anoxic realms of a high saline aqua layer separated by a less saline upper stratified aqua column (Ten Haven et al. 1988). Gammacerane is primarily used to calculate salinity in stratified water columns and is thought to be a marker of anoxic photic zones (Adebanji et al. 2015). The investigated G30 vs C30 ratio is between 0.05 and 0.24 (Table 2), indicating the presence of saline in the stratified aqueous column via the sedimentation of lower and upper post-rift deposits in the Orange Basin. Furthermore, G30/C30 indicates that the paleoxygenation condition of sediment accumulation was decreasing, as supported by Pr/Ph values of less than 0.8, even though the source rocks have freshwater lacustrine influences (Fig. 4c).
Relative sterane richness is a significant quality shift in OM’s biological origin that distinguishes source distinctness while distributing cholesterol from organisms (Fig. 5b, c). C29 sterane is abundant in terrestrial land plants, whereas animals and aquatic plants are primarily composed of sterols of C27 molecular weight, followed by C28 sterols. Sterane C27 is primarily found in algae, while Sterane C29 is found in higher terrestrial plants. Limnic conditions are characterized by low C28 sterol concentrations (Volkman 1986). Similarly, micro-algae and cyanobacteria are good sources of C29 sterols. The analyzed shales contain more C29 sterols (22.00–63.77 percent) than C27 (15.22–44.00 percent) and C28 (18.48–36.46 percent) steranes, implying abundant terrestrial input as opposed to planktonic bacterial to algae organic matter, as shown by the regular steranes triangular plot (Fig. 5b). In other words, except for KH1-3510, which is “L” type, the regular sterane is an anti-“L” type that is predominantly found in land plants (C27/C29 < 1). It should be noted that the dominance of C29 in planktonic land plant input is relatively much higher in KE1-1250 (63.77 percent) and in KH1-3510 (22 percent). A tri-plot of the C27, C28, and C29 distributions indicates the type of primal matter as mixed marine–terrigenous source inputs (Fig. 5b, c). The relatively low sterane vs hopane ratio (Table 3) distinguishes terrestrial or microbial rework matters (Tissot and Welte 1984). Steroids and hopanoid constituents differ between organisms; differences in sterane versus hopane allow only a qualitative evaluation of eukaryotes and prokaryotes’ source inputs (Peters et al. 2005; Seifert and Moldowan 1978). Pr/Ph and C27/(C27 + C29) R values range from 0.10–0.58 to 0.19–0.67, respectively, indicating that anoxic primal matter predominates in terrestrial/coastal anoxic environments (Fig. 5d). Furthermore, most samples have a highly anoxic paleoxygenation condition, with Pr/Ph values less than or equal to 0.58 (Table 1).
According to Phenanthrenes, the source input index also provides information on the origin of marine organic matter (Fig. 6a). The Paq index supports this claim, which is defined as the absolute input of aquatic macrophytes as well as higher terrigenous remains (Ficken et al. 2000). Paq values less than 0.1 indicate unmixed terrigenous matter, while Paq values between 0.1 and 0.4 indicate abundant emergent plant contribution, and Paq values between 0.4 and 1 indicate abundant contributions of emergent macrophytes and submerged plant remains under anoxic conditions (Fig. 6c). There is a high concentration of marine inputs as indicated by Paq from 0.7 to 0.94 (Table 1). Furthermore, a measure of methyldibenzothiophene isomers (MDBT) in the Orange basin source rocks varies as follows: 1 ˃ 4 ˃ 2 + 3 methyldibenzothiophene (Fig. 6b) synonymous to carbonate to shaly organofacies (Hughes et al. 1995). In addition, the peak distribution of methylphenanthrenes (9˃1˃2˃3; Fig. 6a (i)) and the predominance of 9-methylphenanthrene (9MP) over its isomers suggests that organofacies originated from marine input (Budzinski et al. 1995; Escobar et al. 2011). Most samples contain dimethylphenanthrenes and dibenzothiophene dimethylated duplicates. Based on aromatic biomarker analyses, the studied source rocks were accumulated in reduced anoxic carbonate settings. In conclusion, biomarker analysis reveals that the well “K” sediments studied are genetically derived from the coastal marine environment as Pr vs Ph values ranging from 0.10 to 0.58 support this, indicating marine anoxic conditions (Fig. 3b; Taohua et al. 2018). These source rocks contain a significant number of algae and microbial marine primal matter, as well as some mixed organic remains deposited in a more reducing environment (Fig. 3c).
Maturation of organic matter
This expands on the extent to which OM is intensely heated and transformed into petroleum. The molecular specification for the ratio and distribution of biomarkers has been modified to investigate the degree of thermal maturation (Gogoi et al. 2021). To investigate the thermal maturation of the shaly sediments, a few molecular maturation determinants, such as the C29ββ(ββ + αα) ratio, the C3222S/(22S + 22R) ratio, and the M30/H30 plus methylphenanthrene indices (MPI), were adapted (Table 1).
The sterane isomerization model is widely accepted as an OM maturity guide (Seifert and Moldowan 1978; Xuijian et al. 2016). C29 αααR/ (αββR + 29αααR) ratios in examined shales range from 0.41 to 0.74 (Table 3), a ratio of C29αααS/(20S + 20R) varies from 0.10 to 0.51. C32 22S/(22S + 22R) hopane ratios (0.44–0.82) demonstrated that an equilibrium had been reached and that the shales were mature for hydrocarbon generation (Seifert and Moldowan 1986). With increasing thermic maturation, these fractions increase (Adekola et al. 2012). In the source rocks, the maximum C29αααS/(20S + 20R) values are 0.51 or less. These values propose that some of the examined shales are thermally mature to generate hydrocarbons, while others are immature (Fig. 5e). The mortane versus hopane (M30/H30) reduces with rising thermic maturation between 0.8 for the immature extract and 0.15–0.05 matured sediments (Mackenzie et al. 1980; Peters et al. 2005). M30/H30 examined is from 0.14 to 1.00. KD1-1900 is immature, with an M30/H30 ratio of no more than 1.00.
C29diasterane versus regular steranes were also suitable for detecting OM thermal maturity. Diasteranes are formed by the transformation of sterols, and diagenesis, as well as catagenesis, making them more balanced than steranes at higher maturation (Peters and Moldowan 1993; Peters et al. 2005). C29diasterane/29αααR ranges from 0.09 to 1.18, indicating that most shales are mature enough for hydrocarbon generation. Other biomarker maturity parameters are Ts vs (Ts + Tm), C3222S/ (22S + 22R), M30/H30, the 20S/ (20S + 20R), and C29ββ/(ββ + αα) (Mackenzie et al. 1980; Waples & Machihara 1991). Source input and thermal maturity of OM are represented by these ratios. This is based on OM source input and maturity in the case of Ts vs (TS + Tm) and Ts vs Tm (Peters and Moldowan 1993). Tm is unstable during the catagenesis of hydrocarbon generations; as a result, Ts versus Tm increases with maturity (Seifert and Moldowan 1978). In the analyzed shales, the Ts/ (Ts + Tm) is from 0.09 to 0.57 (immature to early maturity).
The Ph/n-C18 and Pr/n-C17 are homogeneous reflecting similar interpretations as for low-to-high thermal maturity (Fig. 3c). In comparison with other biomarker maturity parameters (Fig. 5e) and vitrinite reflectance measurements, the methylphenanthrenes index (Table 1; Radke and Welte 1983; Radke et al. 1982) yielded computed reflectance of vitrinite (VRc percent) ranging from 0.58 to 1.39. (Table 1). Despite inconsistencies in somethermal maturity parameters, the general trend in sample maturity is that, it increases with depth across each respective well. This disparity in maturation is attributed to the Orange basin’s intricate infilling evolution and mixing.
Elemental concentrations and distributions
The genesis, conservation, accumulation, and paleoenvironmental setting of organic materials are significantly influenced by their bulk inorganic composition (Mohialdeen and Raza 2013; Shu et al. 2013; Jia et al. 2013; Moosavirada et al. 2011). Tables 4 and 5 summarize the inorganic geochemical outcome within the examined shales, which includes major and trace oxides with applicable geochemical ratios.
Major elements are commonly found in greater than 1% of major ores. These oxides were investigated using ICP-MS and XRF equipment (Roy and Roser 2013; Peters et al. 2005). SiO2, Al2O3, Fe2O3, CaO, MgO, K20, Na20, SO3, TiO2 have average compositions of 58.20 percent, 15.65 percent, 10.10 percent, 3.92 percent, 2.50 percent, 2.44 percent, 2.33 percent, 1.83 percent, and 1.04 percent, respectively. Less than 1 percent is the average concentration of other major oxides, such as Rb2O, Zr2O, Sr2O5, P2O5, BaO, CuO, and MnO. The major oxide have high SiO2 and Al2O3, indicating a quartz and clay mineral predominance, as expected in siliciclastic deposits. Al/Si ratios ranging from 0.242 to 0.293 are extremely low, with an average of 0.269, implying that Si comes from sources other than clay minerals (Fu et al. 2011). High SiO2 and Al2O3 elements demonstrate detrital matter source input (Ross and Bustin 2009), which is supported by titanium (Ti) oxide at a level of about 1.04 (Table 4) because Ti coexists within terrestrial remains as well as coarse-grained matter influenced under higher energy conditions (Ross and Bustin 2009; Calvert et al. 1996). In the studied shales, iron oxide Fe2O3 is the third most prevalent element (averaging 10.10wt%), and it is usually found in association with Fe sulfides (pyrites and marcasite). Its abundance was detected under petrographic observation as pyrite. The major oxide bivariate scatter plots by Herron (1988) are shown in Fig. 9c; this geochemically classified the investigated Orange Basin samples as shales as demonstrated by the various well lithologic descriptions. The average CaO concentration of 3.92 indicates a calcium-rich environment of deposition (marine setting) toward the depocenter of the Orange basin. Phosphorus, another type of major oxide, was also discovered within the block K, Cretaceous shales (Table 4). The observed shales have a moderate phosphorus amount ranging from 0.127 to 0.378, implying a source of organic matter that is mixed and a reducing sedimentation environment (Shu et al. 2013). Paleoxygenation conditions as determined by biomarker distribution also support this (Table 1). Phosphorus (P) detection in sediments indicates some level of nutrient derivation for biological production (algal growth) under reducing conditions that favor organic matter conservation (Van Cappellen and Ingall 1994; Ingall and Jahnke 1994).
Shales contain trace elements such as Cr, V, Ni, Co, Mo, Mn, and Pb with less than 0.1 percent of each element expressed in parts per billion, ppb (Swaine 1990). The origin of organic matter, paleoxygenation during deposition, organic matter preservation, salinities in water, as well as paleoclimates and tectonic restructuring are determined using major and trace elements (Makeen et al. 2015a; Mohialdeen and Raza 2013; Shu et al. 2013; Jia et al. 2013; Moosavirad et al. 2011; Tribovillard et al. 2005). Cretaceous sediments from the Orange Basin were analyzed using ICP-MS to determine the trace element composition of some post-rift rocks. There are 113,169.94 parts per billion of iron (Fe), followed by aluminum (30,008 ppb), manganese Mn (2,074.73 ppb), and lead (Pb) (830.22 ppb). Among the trace elements found in OM (biophile), the most abundant is chromium Cr, with a concentration of 222.22 parts per billion (ppb) followed by vanadium (221.42 ppb), then nickel (123.53 ppb), cobalt (73.27 ppb), and molybdenum (24.15 ppb) are the least. It is common to analyze source rocks for Ni and V to determine their depositional conditions, oil sources, migration pathways, hydrocarbon maturity, and crude oil classification (Akinlua et al. 2007a, 2007b; Frankenberger et al. 1994; Lewan 1984; Barwise 1990). They are metalloporphyrin complexes that are biophilic in nature incorporated into organic matter or formed during sedimentation (Akinlua et al. 2007a, 2007b; Akinlua et al. 2015). When it comes to identifying oil families, Ni and V are excellent choices (Akinlua et al. 2015; Frankenberger et al. 1994). V/Ni concentrations are similar in all examined shales, and they are genetically related, as shown in Fig. 7a. Vanadium concentrations are lower in KF1-820, KD1-2250, KE1-4105, KH1-1500, and KD1-3120 due to migration. Our findings confirmed that all samples came from the same rock which also affirmed (Akinlua et al. 2018; Fig. 7a).
Paleoenvironmental setting during OM sedimentation
Inorganic geochemical parameters can be used to track physiochemical changes as they occur during sedimentation. Several paleoxygenation sensitive elements, such as V, Ni, Cu, Cr, Mo, and Mn, have been found to be good indicators of paleoredox settings (Adebanji et al. 2014; Mc Manus et al. 1999).
Vanadium (V) vs Nickel (Ni) ratio
The ratios of redox-sensitive elements like V and Ni reveal information about anoxia and deposited matter (Pi et al. 2013; Adebanji et al. 2014). Mostly, sediments contain a comparable amount of these biophile (Lewan 1984). The primary determinants of paleoxygenation conditions during deposition are V and Ni (Galarraga et al. 2008; Bechtel et al. 2001). In this manner, V versus Ni, V versus Cr, and Ni versus Co (Table 5) are sensory indicators of paleoredox settings that differentiate oxic, dysoxic, and anoxic deposits (Galarraga et al. 2008; Ayinla et al. 2017; Hakimi et al. 2015). In a marine environment that ranges from suboxic to anoxic, the amount of vanadium (V) is greater than the amount of nickel (Ni) (Barwise 1990). Vanadium concentrations that are on average higher than nickel concentrations indicate a preponderance of alternating suboxic to anoxic conditions in the area under investigation (Fig. 7b).
Higher V vs Ni ratios indicate marine OM input, while lower V versus Ni ratios indicate terrestrial matters (Akinlua et al. 2018; Galaragga et al. 2008; Akinlua et al. 2007a; Lewan 1984). A greater than one V versus Ni ratio indicates a declining marine environment, whereas a ratio of less than one indicates a terrestrial or mixed source (Akinlua et al. 2015; Vander Spuy et al. 2003).
The proven source rock assemblage of Late Aptian age was determined to be terrestrial, with an average V/Ni of 0.026 (Table 5); this contradicts some reports that Aptian shales are anoxic (Hirsch et al. 2009; Paton et al. 2008; vander Spuy et al. 2003; Herbin et al. 1987). Furthermore, 86 percent of its Cenomanian–Turonian source rocks are marine-derived, while only 14 percent show some mixing (Fig. 7b; Table 5). Co/Ni is adapted to distinguish samples from marine and lacustrine environments in order to differentiate the source rocks into their appropriate origins. The investigated shales have Co/Ni ratios ranging from 0.428 to 1.019 (Table 5). Lacustrine samples have Co/Ni values ranging from 0.79 to 1.07 (Akinlua et al. 2018), accounting for up to 25% of the total rock examined (Table 5). Marine-derived source rocks, on the other hand, have a Co/Ni ratio of less than 0.79, accounting for up to 75% of the sample (Table 5). The V/Ni ratios of some inferred Albian succession source rocks range from 0.641 to 3.768, while Campanian source rocks have V/Ni ratios from 0.981 to 1.873. In general, the investigated shaly samples are primarily of marine origin (Fig. 7a, b), with 75% of the samples being anoxic, and only KF1-820, KH1-1500, KD1-2250, KD1-3120, and KE1-4105 having a V/Ni ratio of less than 1. This suggests that the terrestrial remains in these few samples are insignificant, and that marine organic matter is more prevalent in all the samples.
V/ (V + Ni) fraction
V/ (V + Ni) values greater than 0.5 were found in 75% of the tested samples, indicating reduced bottom water conditions, while 25 percent have values less than 0.5, indicating a dysoxic to anoxic setting of the organic matter (Akinlua et al. 2016). Similarly, V/ (V + Ni) ratios more than 0.84, between 0.54–0.82, and 0.46–0.60, indicates anoxic, eutoxic, and anoxic to dysoxic conditions respectively (Chen et al. 2018; Akinlua et al. 2015). The average V/ (V + Ni) fraction for all investigated shales is 0.56, indicating anoxic to dysoxic and mostly anoxic conditions.
Stable isotope of carbon
Bulk carbon isotope data also distinguishes the varying depositional conditions of organic matter as either marine or mixed matter and depicts the genetic relationship that exists within source rocks (Collisten and Wavrek 1996; Sofer 1984).
According to Sofer (1984), bulk δ13C saturates as well as δ13C aromatic fractions were also utilized to distinguish rock extracts environment (Fig. 8a). δ13C saturate and δ13C aromatic fractions range from -22.29 to -12.24 and -21.38 to -12.4, respectively, with a mean variable 13C signature ranging from −21.84 to −12.32. (Table 6). Source rocks from marine versus terrigenous depositional settings were also distinguished using the canonical variable (CV). CV less than 0.47 denotes marine matters, while CV greater than 0.47 denotes terrestrial plant input (Al-Khafaji et al. 2020; El Diasty et al. 2016; Sofer 1984). The source rocks were derived from the sea, as indicated by the CV range of −8.83 to −2.72 in the extracts under investigation (Table 6). This is consistent with biomarker interpretations and trace elemental properties.
According to Sofer (1984) plots, these δ13C saturated and aromatic fractions were primarily marine-derived (Fig. 8a). As a result, the investigated source rocks are classified as marine source dominated and mixed marine setting with some terrestrial matters. The δ13C of marine organic matter ranges from −22 to slightly less than −20. Mixed marine organic matter following the C4 pathways, on the other hand, ranges from slightly below −12 to up to −18. (Fig. 8a, b). As a result, approximately 57 percent of the samples examined are marine-influenced, while 43 percent are mixed derived via the C4 path. This is also consistent with inorganic geochemistry findings.
Correlation analysis using statistics
For the purpose of describing the geochemical processes that affected the studied shales, the geochemical data was subjected to a Pearson correlation analysis. The Pearson correlation matrix for the investigated shales is shown in Table 7. Except for Al2O3 and K2O, SiO2 having negative correlation with all the major oxides, as shown by the matrix (Fe2O3, CaO, MgO, MnO, Na2O, TiO2). Quartz is the primary source of silica because SiO2 has a negative correlation with all the major oxides (Abubakar et al. 2021, 2019; Sonfack et al. 2021; Ahmad and Chandra 2013). The fact that SiO2 is positively correlated with Al2O3 and K2O indicates that clay mineral is a secondary source of silica (Nagarajan et al. 2017). Al2O3 has a significant negative correlation with Fe2O3, CaO, MgO, MnO, Na2O, TiO2, and P2O5, indicating that it is dissociated from clay minerals (Abubakar et al. 2021). The decomposition of K-feldspars could be a source of kaolin minerals, based on the strong positive relationship between Al2O3 and K2O (Abubakar et al. 2019). Fe2O3 exhibits a strong positive relationship with CaO, MgO, MnO, Na2O, TiO2, and P2O5, implying that these elements are associated with heavy mineral fractions (Sonfack et al. 2021). CaO has a good association with MgO, MnO, Na2O, TiO2, and P2O5, indicating that it is related to carbonate minerals (Nagarajan et al. 2017). In contrast to other major oxides, CaO has a negative correlation, particularly SiO2, Al2O3, Fe2O3, and K2O, signifying that carbonate minerals are the primary source of Ca (calcite), with no influence from detrital sources (Bessa et al. 2021). As indicated by K-feldspar and plagioclase products in the studied shales, chemical weathering is responsible for the negative correlation between both K2O and Na2O (Nagarajan et al. 2017).
Distribution of silica content
The mineralogical content of shales, particularly the calcific content, as well as the quartz content, regulates fracture growth (Ding et al. 2012; Jarvie 2012). These shales range from 51.01 to 63.31 percent silica content (Table 4), with the Upper Turonian KE1-2970 having the highest silica content. Al2O3 has a remarkable range of 12.33–17.85 percent, whereas CaO is generally low, ranging from 1.29 to 11.81%. This shales’ mineralogy is silica dominant, as determined by radiolarian and planktonic tests (Davies 1970). As a result of the presence of reasonable clastic (silica and some carbonates) in the investigated shales, it becomes hard and brittle, allowing it to possibly expel hydrocarbons.
Using the obtained SiO2, geochemical proxies from the examined shales determine the origin of sediments, which can be grouped based on mafic, intermediate, and feldspar content (Armstrong-Altrin 2009). All evaluated samples fall within the compositional field of intermediate igneous rocks, as shown by the plot Al2O3/TiO2 vs SiO2 in Fig. 9a. Similarly, K20/Al2O3 ratios can be used to identify the parent matter of muddy rocks. Samples with a K20/Al2O3 ratio of less than 0.4 show little alkali feldspar, whereas shales with values greater than 0.5 show abundant alkali feldspar in proportion to some other minerals within the shales (Cox et al. 1995). The presence of a small amount of K20/Al2O3 values (0.126–0.191) in the examined shales indicates a low granitic composition in the source area. Furthermore, Al2O3/TiO2 accurately depicts the genesis of source rock (Andersson et al. 2004; Hayashi et al. 1997; Garcia et al. 1994). Whenever Al2O3/TiO2 ratio is 3–8wt percent, it refers to mafic igneous rock, 8–21wt percent denotes an intermediate source origin, and 21–70wt percent denotes feldspar-dominated igneous rock. For the shale samples examined, the Al2O3/TiO2 ratios ranged from 9.998 to 20.311 (Table 4), indicating that the sediments were formed by intermediate igneous rocks (Fig. 9a). It is worth noting that all the computations and illustrative plots performed confirm the origins of Orange Basin shales from Intermediate Igneous rocks, as Zr/Ti plots (Fig. 9b). The major oxide bivariate scatter plots by Herron (1988) are shown in Fig. 9c; geochemically classifying the investigated Orange Basin samples as shales as observed on hand specimen and some lithologic descriptions. Similarly, the major elements discrimination diagram (Fig. 9d) plotted the shales within the intermediate igneous field.
Paleoweathering conditions and sediment maturity
Alkali and alkali earth metal concentrations in rocks determine the degree of weathering in sediments (Nesbitt and Young 1982). Weathering indices are calculated using the following formulas: i) The chemical index of alterations (CIA) and ii) chemical variability index (ICV) as initiated by Nesbitt and Young (1982).
Rb2O/SrO and SiO2/Al2O3 ratios were also calculated to assess the textural and compositional maturity of the Orange Basin shale.
ICV = (K2O + Fe2O3 + Na2O + CaO + MgO + MnO)/Al2O3.
CIA = (Al2O3/[Al2O3 + CaO + K2O + Na2O]) *100 (Nesbitt and Young 1982).
It is worth noting that SiO2 vs Al2O3 ratio of 3 indicates basic igneous rocks, while a ratio of 5 reveals acidic rocks (Cox et al. 1995). In clastic sediments, these ratios are higher than 5, indicating high maturity, while below the SiO2/Al2O3 ratio indicates immaturity (Abubakar et al. 2021; Hernandez-Hinojosa et al. 2018; Cox et al. 1995). The calculated SiO2/Al2O3 ratios range from 3.41 to 4.14, indicating that the investigated sediments are texturally immature (Table 4).
Similarly, an average ICV index greater than 0.85 denotes a typical rock-forming mineral, while less than 0.85 denotes alteration product. Sediment ICV values is from 1.07 to 2.56, with an average of 1.55, signifying weathering activities within the study area (Table 4). Furthermore, ICV versus CIA (Fig. 10) demonstrates that all the original matter analyzed shales plotted above the Post Archean Australia Shales (PAAS) ICV index of about 0.85, indicating immaturity in the composition as well as relatively mild weathering of original source materials (Hernández-Hinojosa et al. 2018).
Another calculation that explains the level of weathering within the source area is the CIA (Nesbitt and Young 1982). Thus, calculated CIA average value is 64.72, oscillating between 45.93 and 74.41 (Table 4), attesting the weathering effect around the parent aggregates (Fig. 10).
Because Sr2+ easily leaches out of terrestrial inputs, the Rb2O/SrO fraction is also used to investigate weathering processes (McLennan et al. 1993). The examined shales have mean Rb+/Sr2+ values ranging from 2.10 to 0.32 above the mean upper continental crust (UCC) value of 0.32, implying some level of weathering activity, as demonstrated by other computations.
Sr2+ and Ba are useful paleosalinity index metals (Wang 1996; Deng and Qian 1993; Liu et al. 1984). They have been significantly modified to depict both seawater and freshwater imprints (Moradi et al. 2016). A high Sr/Ba ratio indicates a high level of salinity, whereas a low Sr/Ba ratio indicates a low level of salinity (Deng and Qian 1993).
Less than 0.20 Sr/Ba ratio indicates freshwater, and more than 0.5 indicates marine water influence (Li et al. 2020; Wei and Algeo 2020). Values ranging from 0.2 to 0.5 indicate a brackish water environment, but caution should be exercised when generalizing the Sr/Ba ratio. The Sr/Ba ratio in the investigated shales ranges between 0.032 and 1.160 (Table 4). In general, 40 percent of the studied shales were derived from seawater, 25 percent from brackish waters, and 35 percent from freshwater settings. The V/Ni vs Sr/Ba cross-plots (Fig. 11b) illustrate relatively low-to-moderate saline stratification as compatible with the varying marine to freshwater influences (0.032–1.160). Due to the low gammacerane index (Fig. 11a; Table 2), stratification during deposition of most samples was minimal signifying low salinity (Peters et al. 2005; Sinninghe Damste et al. 1995).
Paleovegetation and paleoclimatic evolution
Proxy wax (Pwax), proxy aqueous (Paq), and average chain length (ACL) biomarker identifiers are best suited for reconstructing paleovegetation and paleoclimates (He et al. 2019; Ortiz et al. 2013). Pwax deals with the concentration of waxy hydrocarbons derived from higher plants (Zheng et al. 2007). These proxies, specifically n-C23/(n-C27 + n-C31) ratios, measure the comparable amount of Sphagnum species to terrigenous remains to represent the water level within sediments (Andersson et al. 2011). According to He et al. 2019, n-C23/(n-C27 + n-C31) value greater than 0.2 indicates a predominance of sphagnum in a wet environment, as demonstrated by ranges of the analyzed sediments from 0.71 to 5.60. This analysis is supported by Paq values (Table 1), which indicate the contributions of emergent and submerged plants, as well as Pwax values. A humid environment with abundant C27 n-alkanes promotes the growth of woody vegetative cover. Herbaceous remains, on the other hand, are rich in n-C31 and thrive in a dry, warm environment. Even though we are dealing with the dominantly marine-influenced matter as proved by low terrigenous aquatic ratio ranging between 0.07 and 2 (Table 1). The relative number of nC27, nC29, nC31 alkanes is adapted to categorize terrigenous remains as grasses dominated or predominantly from deciduous trees or conifers. Abundant %C31 shows grasses input (Maffei 1996), while the preponderance of %C27 shows more deciduous trees (Oritz et al. 2010). The predominance of conifers is increasing in n-C29 (Schwark et al. 2002). Woody vegetation is associated with a humid climate that contains more C27 alkanes, while herb shrubs thrive in the warm, droughty season (Vegas et al. 2010). A tri-plot of n-alkane %C27, %C29, and %C31 shows a predominance of C27 n-alkanes (Fig. 12), confirming humid paleovegetative climatic cover. It was determined by Lopez-Dias et al. (2013) that woody and herbaceous higher plants contributed to the paleoclimate in different proportions by adapting to the n-C27/n-C31 ratio. n-C27/ n-C31 ratio ranges between 2.00 and 14.00 (Table 1), which corresponds to our more humid climatic phase of deciduous tree growth.
Terrigenous input and paleoclimatic settings are constructed using a variety of major and trace elements (Liang et al. 2018; Li et al. 2020). Paleoclimate reconstruction in the Orange Basin is largely based on C-values. This shows that warm, humid sediments have higher Fe, Mn, Ni, Cr, V, and Co concentration, while arid emplaced deposits are blessed with higher Ca, Mg, Sr, Ba elements. C-value = Σ (Fe + Mn + Ni + Cr + V + Co)/Σ (Ca + Mg + Sr + Ba + K + Na), part per million (Li et al. 2020). This means the greater the C-values, the more humid while lower C-values connotes arid paleoclimate. C-value above 0.8 shows humid nature, 0.6–0.8 signifies a semi-humid climate, and 0.4–0.6 connote semiarid to semi-humid paleoclimatic settings. On the other hand, semiarid paleoclimates have C-values that spans from 0.2 to 0.4 (Cao et al. 2012; Qui et al. 2015), and below 0.2 refers to arid paleoclimates. The shales of the Orange basin have C-values ranging from 0.60 to 1.34 (Table 4) signifying semi-humid to humid paleoclimatic conditions. The Sr/Cu values were adopted to depict the paleoclimatic setting when terrestrial organofacies were emplaced (Liu and Zhou 2007; Moradi et al. 2016; Li et al. 2020). Low Sr/Cu ratios are associated with humid climates, also, the higher the values, the more aridity it is (Liu & Zhou, 2007). Sr vs Cu index ranging from 1.5 to 5 represents warm–humid climate, whereas ratios above 5 connote hot, dry climatic conditions (Han et al. 2020). The investigated shaly sediments are within the warm, humid climatic condition as the Sr/Cu ratio is averagely up to 1.63, ranging from 0.857 to 2.737 (Table 4). The inorganic calculations on the paleoclimate settings of the Orange Basin are therefore tied to the biomarker proxies.
Using a stable isotope of carbon as a determinant of sediment settings and paleoclimate is hard to decrypt. δ13C variability is predominant than observed between macerals of the exact succession; as such, these differences are not from changes in organic matter composition only (Jingjing et al. 2020). A good correlation of maceral indices exists between δ13C and the vegetation index. The vegetation index (VI) computes the maceral constituent ratio (Calder et al. 1991), which is supported by δ13C and other data (Table 6; Fig. 8b).
Therefore, VI = (TEL + FUS + SFU + SUB + RES)/ (DEN + INT + ALG + LPD + SPO + CUT).
TEL = tellocollinites, FUS = fusinite, SFU = semifusinite, SUB = suberinite, RES = resinite, DEN = densinite, INT = inertodetrinite, ALG = alginite, LPD = liptodetrinite, SPO = sporinite, and CUT = cutinite.
Thus, VI (Jingjing et al. 2020) adequately depicts the amount of preserved tissues, degraded tissues, and particulate organic matters present. Almost all less negative δ13C values go with the intermediate VI values (Table 6; Fig. 8b). Higher VI connotes woody vegetation, favoring drier and shallow water table (Calder et al. 1991; Calder 1993; Jingjing et al. 2020). Low VI, on the other hand, may indicate a wetter environment, resulting in larger, more microbially diverse herbaceous plants. Therefore, VI and the mean δ13C association corroborate that there were fluctuating semi-dry and wet intervals in the time of Orange Basin sediment amassment (Table 6).
Tectonic framework restructuring
The tectonic setting in the Orange basin is well-understood via inorganic major and trace element bivariate plots. Plate tectonics influences the rock chemistry (Peters et al. 2005; Makeen et al. 2015a). Using bulk elemental analyses, bivariate plots of K2O/Na2O vs SiO2 were used to reconstruct the tectonic setting of the Orange basin (Roser and Korsch, 1986). The concentration of K2O/Na2O and SiO2 increases from oceanic island arc settings through active continental boundary settings into passive margins. Most of the investigated shales fall within the passive continental margin, with only two samples falling within the active margin (Fig. 13a). Generally, three types of tectonic settings, according to Verma and Amstrong-Altrin (2013): island or continental arc, continental rift, and collisional configuration. Figure 13b depicts plots within the continental rift indicating the Orange basin shales as a byproduct of the rifted passive margin. A similar tectonic episode is thought to have occurred in West and Central Africa at about the same time as the South American plate opened during the Cretaceous (Genik 1992). The opening of the Atlantic Ocean during the Early Cretaceous Period was the primary cause of the separation of the West African plates (Guiraud and Maurin 1992).
The depositional environment, lithotype, age and thermal maturity of Cretaceous post-rift sediments from the Orange Basin were revealed by molecular, bulk geochemical, and elemental characterization. Some of the findings are as follows:
The unimodal pattern of n-alkane, as well as its components such as Pr vs Ph, Pr/nC17 versus Ph/nC18 cross-plot, Pr/Ph vs C29R/C27R, Pr vs Ph, and C27/ (C27 + C29) R, confirms the accumulation of the studied lithofacies in a coastal anoxic environment. To determine the origin of source rock samples, a biomarker with structural similarities to precursor kerogen and biodegradation resistance was used. Based on aliphatic and aromatic compounds, this method revealed that all samples have the same structure. As a result, all the samples were obtained from the same source.
Biomarker finger printing is consistent with inorganic trace elements because samples are primarily from the same source rock from the same family of rocks and derived under anoxic to dysoxic conditions. This is supported by a Pr/Ph ratio of less than one, a homohopane distribution, and a V/ (V + Ni) ratio. Geochemically, the G/C30 value reflects that the source rock must have been accumulated under reduced conditions, with the sedimentary column consisting of a stratified aqueous column.
A wide range of maturation is evident in the samples examined, ranging from immature to late oil generation. Thermal maturity parameters adapted included C3222S/(22S + 22R), C29αααS/(20S + 20R), M30/C30 ratio, C29diasterane/29αααR and the methylphenanthrene index calculated vitrinite reflectance (VRc percent) and measured VRO percent), with 50% of examined samples showing peak oil window (0.676–0.899).
Inorganic trace findings, together with molecular geochemical cross-plots, isotopic δ13C saturate, and δ13C aromatic cross-plots, confirm the marine genetic origins of organic matters with few mixed materials.
The majority of the less negative δ13C values are associated with woody vegetation that thrives in lower standing water tables and less wet climatic conditions. Thus, VI and the mean δ13C association support the existence of fluctuating dry and wet intervals during the time of Orange Basin sediment accumulation.
Source rock inputs were differentiated by adapting Paq, TAR, and total amounts of n-C27, n-C29, and n-C31. A long-chain n-alkane in a variable form, particularly n-C27, demonstrates diversity in vegetation and can be used to reconstruct paleoenvironmental evolution; in the dry season, grasses grow, while trees multiply in the rainy season.
Inorganic computations on the paleoclimatic condition of the Orange Basin agree with biomarker proxies and isotopic studies confirming semi-humid to humid paleoclimatic conditions.
In terms of textural and compositional maturity, shale samples were classified as immature. Tectonic setting discriminant function diagrams showed that the continental rift’s marginal zone was passive. According to this interpretation, the tectonic setting is consistent with the Cretaceous event of West and Central Africa, which separated the South American plate from the West African plate, resulting in the opening of the South Atlantic Ocean and the creation of the Orange Basin.
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Usmanu Danfodiyo University in Sokoto, Nigeria, and the Tertiary Education Trust Fund are thanked for providing the first author with a study leave and a PhD fellowship at the University of Malaya in Malaysia. The authors gratefully acknowledge a research grant from the University of Malaya (GPF019B-2018; IF064-2019). The authors would also like to thank the Petroleum Agency of South Africa (PASA) for supplying the cutting samples. We are grateful to Mr. Zamri and Mr. Aliff for their assistance with the GCMS analysis. The authors gratefully acknowledge Mr. Usman and Miss Aaiinaa’s assistance in crushing some samples, and Dr. Ibraki’s assistance in digesting the ICP-MS samples. The authors would like to thank the editor and anonymous reviewers for providing constructive feedback on an earlier version of the work, which helped to improve it significantly.
The authors declare the following financial interests/personal relationships which may be considered as potential competing interests: Award of PhD scholarship to the first author by Usmanu Danfodiyo University Sokoto through Tertiary Education Trust Fund, Award of research grants by the University of Malaya (GPF019B-2018; IF064-2019).
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Yelwa, N.A., Mustapha, K.A., Opuwari, M. et al. Biomarkers, stable carbon isotope, and trace element distribution of source rocks in the Orange Basin, South Africa: implications for paleoenvironmental reconstruction, provenance, and tectonic setting. J Petrol Explor Prod Technol 12, 307–339 (2022). https://doi.org/10.1007/s13202-021-01317-9
- Trace element
- Stable isotope
- Orange basin
- Tectonic setting