Enhanced Continental Weathering Triggered the Anoxia of Seawater and Mass Extinctions During the Late Ordovician

Abstract

During the Late Ordovician period, changes in climate and mass extinctions were observed; however, the factors influencing these phenomena have not been fully understood. In order to understand the relationships among redox water conditions, climates, and mass extinctions in the Late Ordovician, this study analyzes the chemical index of alteration (CIA) in shales and ⁸⁷Sr/⁸⁶Sr in carbonate leachates as proxies of changes in chemical weathering intensity and chemical weathering rate in the Late Ordovician (mainly from Katian to Hirnantian). The results show that an enhanced chemical weathering rate (increased ⁸⁷Sr/⁸⁶Sr ratios) and decreased chemical weathering intensity (decreased CIA values) characterized the late Katian, which might be attributed to the global orogenesis and enhanced precipitation/runoff under the warming climate (late-Boda warming). This enhanced chemical weathering rate contributed to the CO₂ drawdown in the P. pacificus biozone, corresponding to the initiation of cooling and further glaciation. Meanwhile, the enhanced weathering-induced high primary productivity could have contributed to the expansion of anoxic seawater in the Katian, which further caused the Katian extinction. The Hirnantian Glaciation was characterized by high ⁸⁷Sr/⁸⁶Sr ratios in carbonates and extremely low CIA values in shales, which were likely related to the exposure of continents during low sea level and the glacial grinding of unweathered rocks. This study shows that the highest denudation rate and lowest chemical weathering intensity in the Hirnantian stage might have resulted in enhanced CO₂ release and contributed to the end of glaciation.

1. Introduction

The Late Ordovician witnessed significant geological events involving climate changes, mass extinction, and biological recovery [1,2,3,4]. The former includes the warming of the climate (Boda warming event in late Katian) and the cooling of the climate (glaciation in Hirnantian), and the latter generally includes the Great Ordovician Biodiversification Event (GOBE) and Late Ordovician mass extinctions (LOMEs). Extensive studies have shown that Boda warming and Hirnantian Glaciation were related to changes in atmospheric CO₂ and temperature caused by a combination of intense volcanic eruption, increases in continental weathering, and organic carbon burial [5,6,7,8,9]. Two phases of extinction marked the LOMEs in traditional understandings, and new studies have found that the LOMEs happened within the middle–late Katian (named as Katian extinction) [10,11]. A variety of extinction mechanisms have been proposed, mostly involving temperature, volcanism, and euxinia in seawater [12,13,14,15,16,17].

The chemical weathering of silicate rock entails the dissolution of silicate minerals and the formation of secondary minerals [18]. This process transforms atmospheric CO₂ into dissolved HCO₃⁻ and releases significant amounts of nutrients (such as phosphorus and iron) into the ocean, thereby substantially influencing the paleoclimate, primary productivity, and oceanic hydrochemistry conditions [19,20,21]. Factors such as surface temperature, tectonic activity, volcanic eruptions, and land-plant colonization generally impact the rate of chemical weathering [22,23,24]. It is suggested that the intense tectonic activity during the Late Ordovician not only led to crustal uplift [5] but also facilitated the southward movement of the Appalachian orogenic belt relative to the Laurentian continent [8], both of which enhanced continental weathering. Moreover, the greenhouse climate resulting from significant volcanic activity in the late Katian and a plentiful supply of weathered protoliths further amplified continental weathering [25,26]. In contrast, the global cooling in the Hirnantian was characterized by weak chemical weathering intensity [27].

Silicate weathering mainly includes chemical and physical denudation rates (denudation rate = chemical weathering rate + physical weathering rate), with chemical weathering intensity representing the ratio of the chemical weathering rate to the denudation rate [28]. Generally, a linear array on the logarithmic plot of chemical weathering versus physical erosion rate could be observed under a low erosion rate (transport-limited), in which case chemical weathering intensity is enhanced due to thick soil (completely weathered) sequences [28]. In contrast, silicate weathering shows substantial scatter away from the trend at higher erosion rates (kinetically limited) when chemical weathering rates are higher [28]. Traditionally, weathering-controlled changes in climate mean that an enhanced chemical weathering rate, rather than increased chemical weathering intensity, contributes to the drawdown of atmospheric CO₂ [18,19,20]. The relationship between climate and weathering is complicated, and several hypotheses have been proposed. For example, a high chemical weathering intensity triggered CO₂ drawdown following a carbon release event caused by volcanism [29], a low chemical weathering intensity and high chemical weathering rate occurred under a warming climate [30,31], and the silicate weathering rate reached a maximum for intermediate denudation rates and further contributed to CO₂ drawdown. In contrast, extremely low denudation rates (supply limited regime) and high denudation rates (kinetically limited regime) can decrease or even reverse CO₂ sequestration [32].

The strontium (Sr) isotopic compositions in seawater are influenced by continental runoff and mantle contributions from hydrothermal systems [33,34,35,36,37,38]. As a result, seawater ⁸⁷Sr/⁸⁶Sr ratios could be used as significant indicators for continental weathering trends in geological periods because these ⁸⁷Sr/⁸⁶Sr ratios could be used to constrain continental weathering fluxes of dissolved Sr [33]. Weathering proxies in mudstones include the chemical alteration index (CIA), plagioclase alteration index (PIA), and chemical weathering index (CIW) [39,40,41]. All of these weathering proxies should reflect the changes in chemical weathering intensity. In south China, the redox water conditions, hydrodynamic conditions, volcanism, and evolution of biomass have been widely analyzed in previous studies [3,10,16,29]. Consequently, detailed investigations into the weathering trends (chemical weathering rate and intensity) of the Late Ordovician (using weathering indices in shales and ⁸⁷Sr/⁸⁶Sr in carbonate leachates) are essential to elucidate the critical relationships between continental weathering, paleoclimate, and the LOMEs.

2. Geological Setting

South China was located as a separate continent in the Ordovician (Figure 1A) [42,43], consisting of the Yangtze Block in northwest and the Cathaysia Block in the southeast. The collision of them happened in the early Neoproterozoic (Figure 1B), leading to the compression between the Yangtze Block and Cathaysia Block [44,45]. The graptolite in the Katian includes Dicellograptus complanatus, Dicellograptus complexus, Paraothograptus pacificus biozones, Metabolograptus extraodinarius, and Metabolograptus persculptus biozones in the Hirnantian and Akidograptus ascensus, Parakidograptus acuminatus, Cystograptus vesiculosus, and Coronograptus cyphus biozones in the Rhuddanian (Figure 1C) [46,47].

A shallow water section (Wuke section) was located in western Sichuan Province, south China (Figure 1B) [17]. The stratigraphy of the Wuke (WK) section is organized into the Linxiang (LX) Formation, the Tiezufeike (TZFK) Formation, and the Butuo (BT) Formation, in stratigraphic order. The LX Formation (~5 m) is primarily characterized by argillaceous limestones and calcareous shales. The Tiezufeike Formation (~15 m) is characterized predominantly by limestones and dolomitic limestones, with the upper part of Tiezufeike Formation containing a significant presence of shelly fauna, notably the Hirnantian fauna. The Butuo Formation (~>20 m) mainly consists of laminated limestones and calcareous mudstones, alternating with layers of argillaceous siltstones and mudstones.

The two boreholes (YD1 and DY3) analyzed in this study were located on the Upper Yangtze platform (Figure 1B). The boreholes are categorized into the Lingxiang Formation, Wufeng Formation, Guanyinqiao Bed, and Longmaxi Formation, arranged in stratigraphic order. The Wufeng and Longmaxi formations primarily consist of black shale, whereas the Guanyinqiao Bed comprises carbonaceous and muddy limestones containing abundant benthic fauna [16].

Environmental differences were noted between the western margin of the Yangtze Block, characterized by a deep shelf (shale dominated), and the southwestern Yangtze Block, characterized by a shallow shelf (carbonates dominated). In the deep shelf area, a widespread distribution of organic-rich shales of the Wufeng (WF) and Longmaxi (LMX) formations were observed [13,16,17]. The siliceous and argillaceous shales of the Wufeng and Longmaxi formations, dating from the latest Hirnantian to Aeronian, were deposited under anoxic water conditions [13,16,17]. The controlling mechanisms for this included intense upwelling, volcanism, and thermohaline stratification. In the shallow shelf areas, minor carbonate facies were present along the flanks of the Kangdian and Central Guizhou uplifts, where the waters likely became less anoxic and upwelling weakened compared to those in the deep-water shelf [41,43]. Also, the carbonates in the shallow shelf were deposited under more oxic water conditions, leading to less total organic matter in the rocks [17].

3. Methods

In this study, a total of 15 carbonate samples from the Wuke section and 29 shale samples from DY3 and YD1 boreholes were performed for geochemical analyses.

3.1. Major and Trace Elements

The melting method was used for the pretreatment of samples for major element analyses, with the cosolvent being lithium fluoride, lithium metaborate, and lithium tetraborate and the oxidant being ammonium nitrate and lithium bromide. The melting temperature was set to 1050 °C, and the melting time was 15 min. An X-ray fluorescence spectrometer (XRF) was used for the analysis of major elements, with the standard curve using Chinese National Standard Material (GBW07101-14) [48] and a relative standard deviation (RSD) of less than 2% [48]. The analytical precision was better than 10%.

The trace elements were analyzed by an Agilent 7700e ICP-MS at the Wuhan Sample Solution Analytical Technology Co., Ltd., Wuhan, China. First, fresh and fine-grained shale samples were powdered into 0.075 mm-diameter pellets, and these powders were placed in an oven at 105 °C for approximately 12 h. Second, approximately 50 mg of the sample was dissolved in HNO₃ and HF. Third, HNO₃, Milli-Q water, and an internal standard solution were added. Fourth, the final solution was transferred to a polyethylene bottle and diluted to 100 g by adding 2% HNO₃. The analytical precision was better than 5%.

The CIA is calculated as follows: CIA = Al₂O₃/(Al₂O₃ + CaO* + Na₂O + K₂O) × 100 [49]. In the equation, CaO* represents the CaO content in the silicate, and thus, it is calculated following the method: CaO* is initially corrected using P₂O₅ contents (CaO − 10/3 × P₂O₅) [50]. If the calculated value exceeds the Na₂O content, the CaO* value is set as the Na₂O value; otherwise, the CaO* value is set as the CaO content. Due to the possible addition of K₂O to weathered samples during K-metasomatism, a correction of CIA values (CIA_corr) has been proposed as follows: CIA_corr = Al₂O₃/(Al₂O₃ + CaO* + Na₂O + K₂O*) × 100 [51], where K₂O^* is calculated as follows: (K₂O*) = {m × [(Al₂O₃)] + m × [(CaO*) + (Na₂O)]}/(1 − m), with m = (K₂O)/[(Al₂O₃) + (CaO^*) + (Na₂O) + (K₂O)].

To eliminate the influence of K₂O, Harnois [39] proposed the CIW proxy (CIW = Al₂O₃/(Al₂O₃ + CaO* +Na₂O) × 100). Fedo et al. [51] defined the PIA proxy as a necessary weathering intensity, which is calculated as follows: PIA = (Al₂O₃ − K₂O)/(Al₂O₃ − K₂O + CaO* + Na₂O) × 100. In addition, the compositional maturity of a rock (ICV) is expressed as ICV = (Fe₂O₃ + K₂O + Na₂O + CaO + MgO + MnO + TiO₂)/A1₂O₃ [52].

3.2. Elemental Concentrations and ⁸⁷Sr/⁸⁶Sr Analysis in Carbonates

About 100 mg of carbonate samples were rinsed three times with the Milli-Q water (18.2 MΩ). The carbonate samples were leached twice with 0.05 M acetic acid at room temperature for 24 h. Fractions of the carbonate leachates were retained in 2% HNO₃ for elemental analysis using an Elan Quadrupole ICP-OES. The reference material JLs-1was analyzed [53] and indicates that accuracy and precision were less than ±7% for all elemental concentrations reported here.

Sr isotopic ratios were measured using the Neptune plus MC-ICP-MS at the Wuhan Sample Solution Analytical Technology Co., Ltd., Wuhan, China. The Sr isotope analysis of carbonate leachates involves three steps [54]. Sample digestion: about 50–200 mg of sample power (200 mesh) was placed in an oven at 105 °C for drying and was then dissolved using the HNO₃ and HF. The solution was evaporated to dryness and was dissolved in 1.0 mL of 2.5 M HCl again. Column chemistry: the supernatant solution was loaded into an ion-exchange column packed with AG50W resin. The Sr fraction was eluted using 2.5 M HCl and gently evaporated to dryness prior to mass-spectrometric measurement. Sr isotope analyses were conducted at the Wuhan Sample Solution Analytical Technology Co., Ltd., Hubei, China, using a Neptune Plus MC-ICP-MS. International NIST 987 standard [54] was measured every seven samples analyzed, with the ⁸⁷Sr/⁸⁶Sr ratio of 0.710242 ± 14 (2SD, n = 345) having been achieved, which is identical within error to their published values 0.710248 ± 12 [55]. Additionally, the BCR-2 (basalt) and RGM-2 (rhyolite) yielded results of 0.705012 ± 22 (2SD, n = 63) and 0.704173 ± 20 (2SD, n = 20) for ⁸⁷Sr/⁸⁶Sr, respectively.

4. Results

4.1. Major Elements

The result of major element concentrations can be found in the Supplementary Files (Tables S1 and S2). Si, Al, and Ca are the primary elements and show high concentrations in the studied shales, while Fe, Mg, Na, K, Ti, and P are the minor elements and show low concentrations (Tables S1 and S2). Generally, Si, Al, Na, and K are linked with silicate minerals, and Ca and Mg are linked with carbonate minerals. The Si percentage varies from 35.7% to 70.0% in DY3 (except for one marlstone with low SiO₂ contents (13.4%) in Guanyinqiao Member) and from 56.0% to 77.7% in YD1, Al varies from 3.2% to 11.5% in DY3 and from 4.7% to 15.7% in YD1, and Ca varies from 1.7% to 26.0% in DY3 and from 0.6% to 15.7% in YD1.

All the CIA, CIA_corr, CIW, and PIA profiles in DY3 and YD1 boreholes exhibit decreasing trends from the Katian to Hirnantian (Figure 2), with values ranging from 58.1 to 66.0, from 59.2 to 67.9, from 69.5 to 81.0, and from 62.1 to 75.4 in DY3 and from 60.4 to 68.3, from 61.1 to 70.5, from 69.3 to 83.5, and from 64.0 to 78.4 in YD1. The ICV shows variable values throughout the Katian to Hirnantian rocks, ranging from 0.98 to 14.4 in DY3 and from 0.90 to 4.43 in YD1.

4.2. Trace Elements

The trace elements in shales of DY3 borehole have been analyzed in this study (Table S1). Sc, Hf, Rb, La, and Th in the Katian show steady contents, varying from 8.40 μg/g to 11.2 μg/g, from 2.51 μg/g to 4.01 μg/g, from 104 μg/g to 148 μg/g, from 31.8 μg/g to 47.6 μg/g, and from 12.5 μg/g to 17.8 μg/g (Table S1). The Sc, Hf, Rb, Th in these rocks show relatively low contents compared to those in PAAS (Sc: 16, Hf: 5, Rb: 160), while the La and Th show more variables when compared with those in PAAS (La: 38, Th:14.6). Sc, Co, Rb, La, Th in the Hirnantian show variable contents, varying from 3.60 μg/g to 9.40 μg/g, from 1.59 μg/g to 4.17 μg/g, from 38.2 μg/g to 107 μg/g, from 15.5 μg/g to 69.2 μg/g, and from 4.85 μg/g to 13.8 μg/g (Table S1). Except for La, all other trace elements show lower contents relative to those in PAAS.

4.3. ⁸⁷Sr/⁸⁶Sr and Elements in Carbonate Leachates

The results of ⁸⁷Sr/⁸⁶Sr and elements in carbonate leachates can be found in the Supplementary Files (Table S3). ⁸⁷Sr/⁸⁶Sr have stratigraphic variations in the WK section, increasing from 0.708492 to 0.709022 in the lower Tiezufeike Formation and decreasing from 0.709022 to 0.708130 in the upper Tiezufeike Formation. In addition, the Sr and Mn in carbonate leachates show variable contents throughout the Tiezufeike Formation, ranging from 301.1 μg/g to 1446.6 μg/g, and ranging from 17.0 μg/g to 183.7 μg/g, respectively. The corresponding Mn/Sr and Sr/Ca ratios vary between 0.03 ppm/% and 0.50 ppm/% and between 29.91 ppm/% and 190.68 ppm/%, respectively.

5. Discussion

5.1. δ¹³ C Records and Climate Events in the Late Ordovician

Carbon isotopes are identified as a useful tool for determining local and global stratigraphic correlations. We compared the δ¹³C values in Late Ordovician sediments of the WK section [58] with those in other published studies [59,60,61,62]. In south China, the δ¹³C_org values are high in the D. complanatus Biozone (Figure 3; Linxiang Formation), and they shift toward low values within the D. complexus Biozones (lower Wufeng Formation in shale-dominated sections and lower Tiezufeike Formation in carbonate-dominated section). The δ¹³C_org values remained low throughout the lower-middle P. pacificus biozone; subsequently, the δ¹³C_org values began to increase slowly in the upper P. pacificus biozone and more rapidly in the lower Hirnantian (P. persculptus) zone and peaked in the upper part of the M. extraordinarius biozone. This peak in high δ¹³C_org (~−25‰) is concurrent with the occurrence of the Hirnantia–Dalmanitina fauna [58], predominantly a cold/cool-water fauna found in the upper Tiezufeike Formation in WK section and Guanyinqiao Member in the DY and YD boreholes (Figure 3).

Various climatic events, including the early and late-Boda warming, mid-Boda cooling, and Hirnantian Glaciation, have been delineated through petrological and geochemical evidence, with δ¹³C record correlation utilized to constrain the timing of these geological phenomena (Figure 3; [63]). δ¹³C values remained stable and high during the late Katian (from D. complanatus to D. complexus), aligning with the mid-Boda cooling [61]. Subsequently, δ¹³C values decrease in the P. pacificus biozone, reflecting intense volcanism and late-Boda warming [64]. An increase in δ¹³C is noted in the uppermost P. pacificus Biozone, equivalent to the upper Tiezufeike Formation in WK, the upper Wufeng Formation in Tianjiawan and Qiliao, and the Vinini Formation in Vinini Creek (Figure 3). A positive δ¹³C excursion is observed within the M. persculptus and M. persculptus biozones, alongside a significant regression and the most extensive Hirnantian glacial event. By the onset of the Silurian, δ¹³C values revert to lower levels similar to those before the Hirnantian Glaciation, marking the end of the glaciation and a transition to a warmer climate (Figure 3).

5.2. The Reliability of the Weathering Proxies in Shales and Carbonates

5.2.1. Evaluation of Weathering Proxies in Shales

The weathering proxies in shales are widely utilized to evaluate silicate weathering [65]. In this study, CIA, CIA_corr, CIW, and PIA show decreasing trends throughout the DY3 and YD1 boreholes (Figure 3; [66,67]) and positive correlations among these proxies could be observed (Figure 4A–C; [68,69]). A thorough evaluation of these proxies is essential because they might be affected by source rocks, hydrodynamic cycles, and post-depositional diagenesis (Figure 4A–F). In this case, this study uses the trace element concentrations and their ratios, Al₂O₃/SiO₂, and K/Rb, to evaluate the influences of the above factors.

Traditionally, Zr, Ti, Al, and their ratios are used to identify the lithology of parental rocks; however, Zr is strongly controlled by the presence of zircons and hence the sedimentary rocks, as zircons tend to be enriched in the silt-size fraction. In addition, Ti abundances in shales generally co-vary with Al, because both elements are immobile during chemical weathering and associated with secondary clay minerals [68]. So, the granulometry of shales is likely the dominant parameter controlling Al₂O₃/TiO₂ ratios in shales. Instead, we choose elements such as Hf, La, Th, and Sc in clastic rocks as effective indicators for determining the types of source rocks and tectonic settings [69]. The discrimination diagrams indicate that all the samples are near the regions of felsic to intermediate rocks (Figure 5A,B; [70,71,72]).

Additionally, A-CN-K diagrams are utilized to determine the types of source rocks. The deviation of our samples from the ideal weathering trend of source rocks (parallel to the A-CN axis) in the Al₂O₃− (CaO* + Na₂O)− K₂O (A-CN-K) ternary diagram (Figure 4F). The corrected CIA values (CIA_corr) are distributed along the ideal weathering trends of source rocks, and the intersection between the weathering trend line and feldspar join represents the source rock. In this study, the ternary diagrams show that the samples from two boreholes converge towards a point representing the source rocks (granodiorite or granite; Figure 4F). These comprehensive observations strongly support the conclusion that Late Ordovician shales derived from a consistent protolith source. Therefore, the observed variations in weathering indicators throughout the Katian and Hirnantian periods are unlikely to be attributable to changes in the types of source rocks.

Adding K₂O during diagenesis can affect the accuracy of the CIA values because the samples are characterized by less weathering than their actual state. The K/Rb ratio can be used to assess the degree of K mobility due to metasomatism, with the rocks having lower K/Rb ratios (>220) in potassium metasomatism [66]. We found that our samples are characterized by slightly high K/Rb ratios (200–227), indicating minor influences of diagenesis (Figure 3). The CIA_corr values are calculated to avoid the influence of potassium metasomatism. In our study, a substantial positive correlation between CIA_corr and CIA indicates that the correction of the CIA preserves the temporal weathering trend (Figure 4A). Furthermore, the variations in weathering proxies might be related to multiple sedimentary cycles before burial, thus influencing the actual weathering experience of the protolith in source areas. Generally, the ICV is a reliable indicator for distinguishing mature and immature mudstones, with mature mudstones experiencing several weathering cycles and exhibiting low ICVs (<1). In contrast, immature mudstones exhibit high ICVs > 1 [67]. In our study, all the shales have high ICVs > 1, with just three samples showing slightly low values (0.9, 0.92, 0.98), suggesting that they experienced only one cycle of weathering (Figure 4D). The Al/Si ratio is a grain size proxy, with fine particles having high Al₂O₃/SiO₂ ratios and coarse components having low Al/Si ratios. There is moderate correlation between CIA_corr and Al₂O₃/SiO₂ in YD1, which could not be observed in DY3 (Figure 4E). We believe that there is a negligible influence of grain size on the variations in weathering proxies for the DY3. In YD1, the coarse part is characterized by low chemical weathering intensity, likely enhanced erosion of unweathered rocks with high size.

5.2.2. Evaluation of Carbonate ⁸⁷Sr/⁸⁶Sr

The Sr isotopic compositions in carbonates are susceptible to early diagenetic transformations and late diagenetic fluid reactions [73]; therefore, evaluating the reliability of carbonate ⁸⁷Sr/⁸⁶Sr before using it as the primary seawater record is important [74]. The seawater Sr is incorporated into carbonate minerals in the synsedimentary stage; however, the Sr concentrations in carbonate minerals could be modified during diagenesis [75]. Generally, 300 ppm Sr appears to be the lower threshold value, in which case the bulk carbonate could record the primary ⁸⁷Sr/⁸⁶Sr ratios [37]. All of the studied samples have high Sr contents (Figure 6) and there is no obvious relationship between Sr and ⁸⁷Sr/⁸⁶Sr (Figure 6), suggesting that primary seawater ⁸⁷Sr/⁸⁶Sr is potentially preserved in bulk carbonate.

Mn tends to be incorporated into carbonate minerals when bulk carbonate reacts with nonmarine diagenetic fluids, leading to higher Mn concentrations and radiogenic ⁸⁷Sr/⁸⁶Sr ratios in bulk carbonate [74]. As a result, Mn/Sr ratios are often used to discriminate diagenetic alterations, with unaltered bulk carbonates characterized by low Mn/Sr ratios [74]. In the WK section, the bulk carbonates have low ratios (Mn/Sr < 1; Figure 6), conforming to the geochemical criteria for Sr isotopes [74,75]. Moreover, bulk carbonate diagenesis should be characterized by low δ¹⁸O values and more radiogenic ⁸⁷Sr/⁸⁶Sr ratios than unaltered bulk carbonate, as diagenesis results in high temperatures. However, there are no correlations between Mn/Sr (Figure 7A), Sr/Ca (Figure 7B), or δ¹⁸O (Figure 7C) and ⁸⁷Sr/⁸⁶Sr, indicating that our samples were not altered by diagenesis.

Dolomitization is a possible process that can alter the geochemical composition (e.g., Mg/Ca, Mn/Sr, and ⁸⁷Sr/⁸⁶Sr ratios) of limestones [76]. However, the peak ⁸⁷Sr/⁸⁶Sr ratios do not correspond to the high dolomite contents in the WK section (Figure 6), and there is no correlation between Mg/Ca and ⁸⁷Sr/⁸⁶Sr (Figure 7D), suggesting that the observed variation in ⁸⁷Sr/⁸⁶Sr is not controlled by primary minerals or dolomitization during diagenesis.

5.3. Fluctuations in Weathering Proxies in the Late Ordovician

5.3.1. Fluctuations in CIA_corr Values in the Late Ordovician

Weathering proxies, including the CIA, CIA_corr, PIA, and CIW, have been used to analyze weathering trends throughout geological history. Previous studies have evaluated the limits of the chemical weathering intensity; however, there are some differences in the threshold values among strong, moderate, and weak [13,45,46]. In this study, the degree of chemical weathering intensity could be divided into intense (~75–100), intermediate (~65–75), and weak (~50–65) chemical weathering intensity [45,46]. It is broadly acknowledged that regions experiencing moderate to high chemical weathering intensity are distinguished by elevated CIA and CIA_corr values. In contrast, low CIA and CIA_corr values are indicative of areas with reduced chemical weathering intensity [13,16].

Within the Yangtze Block, CIA and CIA_corr values demonstrate a progressive decline from the Linxiang Formation to the Wufeng Formation [77,78], despite some variance in absolute values across different boreholes and sections (Figure 8). Similarly, in Scotland, CIA_corr and CIA values also exhibit a marked decrease, with the lower part of the Hartfell shales characterized by higher CIA values and the upper part by lower CIA values (Figure 8; [27]). Consequently, the observed weathering trends based on CIA and CIA_corr in the late Katian are deemed reliable, reflecting consistency across global boreholes and sections (Figure 8). This suggests a diminishing trend in chemical weathering intensity during the late Katian period (from D. complanatus to P. pacificus). Of course, the climatic zonality could influence the chemical weathering intensity, with high latitude areas generally showing low temperatures and low chemical weathering intensity and low latitude areas generally showing high temperatures and high chemical weathering intensity. In this case, the climatic zonality is another possibility influencing the differences in the CIA values between the China and Scotland sections, although this requires more exploration in the future.

In south China, CIA and CIA_corr exhibit lower values in the Hirnantian rocks (CIA average values for the DY and YD boreholes: 60 and 63; CIA_corr average values for the DY and YD boreholes: 61 and 64) than in the late Katian rocks (CIA average values for the DY and YD boreholes: 64 and 66; CIA_corr average values for the DY and YD boreholes: 69 and 67) (

Table 1. The average values of CIA and CIA_corr in the rocks of SD1, DY3, Shuanghe, Qiliao, XY-1, YD1 in south China.

	Borehole/Section	SD1	DY3	Shuanghe	Qiliao	XY-1	YD1
CIA or CIAcorr Value
Average CIA value in Hirnantian		63	60	/	59	59	63
Average CIAcorr value in Hirnantian		69	61	59	/	60	64
Average CIA value in late-Boda warming		66	64	/	66	62	66
Average CIAcorr value late-Boda warming		72	69	65	/	64	67

; Figure 8). This result is consistent with previous studies, suggesting that the rocks deposited during the Hirnantian Glaciation experienced weak chemical weathering intensity [16]. However, increases in CIA and CIA_corr could be observed in Scotland, in contrast to what has been observed in south China. This study attributes these differences in CIA trends to the different weathering regimes under the same climatic conditions. For example, the continental weathering in south China was characterized by the glacial grinding of unweathered rocks with low CIA values; however, a more incongruent weathering regime likely characterized the Scotland rocks based on the high lithium isotopes [27].

5.3.2. Fluctuations in Seawater ⁸⁷Sr/⁸⁶Sr in the Late Ordovician

Several seawater ⁸⁷Sr/⁸⁶Sr records from the Late Ordovician have been reported [79]. The overall published datasets from the mid–late Katian have shown relatively stable and lower seawater ⁸⁷Sr/⁸⁶Sr values (~0.7080), which are attributed to the enhanced weathering of fresh volcanic rocks [80]. However, all previous studies focus more on the Upper Ordovician carbonate successions, especially from the Tremadocian to Sandbian Stages [80]. In fact, there is still the lack of a detailed and high-resolution seawater ⁸⁷Sr/⁸⁶Sr curve from Katian to Hirnantian, which could provide important constraints on the changes in climates.

The overall ⁸⁷Sr/⁸⁶Sr ratios in Late Ordovician section from south China are higher than those in the southwestern Ontario borehole [81] and the Copenhagen Canyon section (Figure 9). However, the initial ⁸⁷Sr/⁸⁶Sr ratios in the Wuke section are similar to the published conodont apatite ⁸⁷Sr/⁸⁶Sr ratios in the late Sandbian and early Katian from the south China sections. It seems that the anomaly of high ⁸⁷Sr/⁸⁶Sr ratios in the late Katian might be caused by a restricted ocean, as is observed in the late Ediacaran West Gondwana basins [82] and the latest Ediacaran Sichuan Basins [83]. This condition of the restricted ocean in the Late Ordovician has also been observed by others, who suggest that moderate restricted conditions characterize the Yangtze shelf sea and likely influence the degree of water mass restriction [84,85]. However, we have re-analyzed the major and trace elements in the WK section to identify if our samples were deposited under a restricted basin [58]. The enrichments of Mo and U in sediments have widely been used to reconstruct the redox state of water columns, as well as the changes in open (seawater) versus enclosed (lake/lagoon) conditions [86]. The results show that a transition from suboxic to sulfidic water conditions could be observed from the early to late Katian, which is followed by suboxic water conditions in the Hirnantian (Figure 5C). Therefore, the majority of samples in our studied section were deposited in open marine settings, instead of restricted ocean. In addition, Co and Mn contents in marine sediments were effective proxies for distinguishing the restricted and upwelling settings due to the significant differences in the supply of these metals to the water column [72]. We found that the majority of samples have stable low Co×Mn values (<0.1), indicating open marine settings or upwelling settings (Figure 5D). As a result, we would like to highlight that the changes in carbonate ⁸⁷Sr/⁸⁶Sr ratios in the Wuke section still record the global trend of seawater ⁸⁷Sr/⁸⁶Sr curve. However, the reasons for the overall high values will be considered in future works.

Several similarities are found in the Late Ordovician ⁸⁷Sr/⁸⁶Sr curves between south China and other areas. First, both the ⁸⁷Sr/⁸⁶Sr curves in south China and southwestern Ontario show low Sr isotopes which is consistent with the overall ratios globally, likely resulting from enhanced continental basalt weathering, a sustainable increase in oceanic hydrothermal Sr flux, or a combination of both [87]. Second, the ⁸⁷Sr/⁸⁶Sr ratios become much more positive in the late Katian, reaching high Sr isotope ratios not attained until the Silurian found in previous studies [80]. Third, the prolonged high ⁸⁷Sr/⁸⁶Sr ratios in the Hirnantian have been observed in south China, southwestern Ontario, and the Copenhagen Canyon (Figure 9). These radiogenic ⁸⁷Sr/⁸⁶Sr ratios in Glaciation were caused by the enhanced mechanical erosion driven by sea-level fall or the preferential weathering of biotite in cooling climates [79]. Therefore, we still believe that our Sr isotope record has global-scale implications for the changes in silicate weathering, hydrothermal activity, and the weathering of fresh volcanic rocks.

The seawater ⁸⁷Sr/⁸⁶Sr is controlled by both river influx from continents and hydrothermal influx [36]. A major drop in marine ⁸⁷Sr/⁸⁶Sr records across the middle–late Ordovician boundary (Darriwilian-Sandbian) was attributed to enhanced continental basalt weathering [88], subduction-related volcanism, or a sustainable increase in oceanic hydrothermal Sr flux [87]. The initial low ⁸⁷Sr/⁸⁶Sr ratios in the lower most Tiezufeike Formation are consistent with the hypotheses about basalt weathering and hydrothermal input; therefore, the influence of these factors should be taken into account. First, A-CN-K diagrams are utilized to determine the types of source rocks, with the results showing that all the samples are sourced from intermediate-felsic igneous rocks, as a result, the effect of changes in the types of source rocks on the ⁸⁷Sr/⁸⁶Sr ratios is minor. Second, the contribution of hydrothermal activity to the deposition of Late Ordovician shales in south China is mainly based on the positive Eu anomaly (Eu/Eu*), because a reduction of Eu³⁺ to Eu²⁺ would occur under the extreme reduction in hydrothermal fluids [89,90]. The results show that weak hydrothermal activity (occasionally slightly positive Eu anomaly) exists on the northern and southwestern margins of the Upper Yangtze Platform, situated in an extensional tectonic background [85]; thus, the possibility of decreasing hydrothermal Sr flux could be excluded. In this case, we just retain the spreading rate or the hydrothermal Sr-input as a steady condition when considering that the spreading rate or the hydrothermal Sr-input is unknown at present. Anyway, we would like conduct more modeling work about the influences of spreading rate or the global hydrothermal Sr-input on the Late Ordovician Sr isotopes in seawater. Third, sea-level-fall-induced restricted seawater conditions are unlikely to lead to the higher ⁸⁷Sr/⁸⁶Sr ratios because the geochemical characteristics in bulk rocks above have identified open ocean settings. Overall, the increase in ⁸⁷Sr/⁸⁶Sr ratios has been attributed to the increasing continental crust erosion during the Taconic orogeny or the sea-level drop related to the onset of Hirnantian Glaciation.

5.4. The Changes in Continental Weathering in the Late Ordovician

Traditional studies have argued that continental weathering is controlled by plant coverage, temperature, runoff, and tectonic processes [91]. First, given that the first non-vascular land plants were only just evolving and colonizing the continents in the mid–late Ordovician, it is probable that the weathering pattern and corresponding clay types were different and less abundant [92]. For example, illites are thought to dominate prior to terrestrialization by plants [92]. In fact, the clay minerals of the Late Ordovician rocks in both south China and Scotland were dominated by illite [27,89]. In addition, the occurrence of non-vascular land plants would be suggested to induce the enhanced formation of clay minerals caused by a decreased denudation rate and increased chemical weathering intensity, leading to the preservation of cations in clays and decreased consumption of atmosphere CO₂ [27]. However, both the changed types of clay and increased chemical weathering intensity were inconsistent with the observation in our geochemical records. As a result, the effect of missing plant coverage on the records of the weathering trends in this study might be minor.

Generally, the elevated temperature could enhance the feldspar dissolution, weathering rate, chemical weathering intensity (high CIA values), and further CO₂ consumption [4]. Hydrologic regulation (precipitation and runoff) has also widely been proposed to control the silicate weathering, with increased runoff causing a high weathering rate in low-relief areas [91]. Some studies further suggested that the runoff-induced high denudation rate would contribute to the high chemical weathering and low chemical weathering intensity [31]. Tectonic uplift can enhance the chemical weathering rate by accelerating the rate of denudation [22]. In this case, high erosion and a rapid supply of fresh minerals would result in a short time to reach weathering equilibrium, so the weathering rate, solute concentrations, and fluxes are high [91].

The weathering flux to the oceans should be maximized when thermodynamic equilibrium between the dissolving and precipitating minerals is approached. Furthermore, peak weathering and CO₂ drawdown occur with moderate erosion rates. In contrast, low and high erosion rates would decrease or reverse CO₂ sequestration. In a simplified model, the low denudation rate would result in a supply limited regime (with thick soils isolating the bedrock from climatic conditions), while the high denudation rate would cause the kinetically limited regime (with rocks being eroded before undergoing significant hydrolysis) [32].

A certain proportion of sedimentary rocks in paleo-catchments will necessarily exaggerate the numerical estimate of the CIA. Therefore, the proposed weathering trends using CIA values of Late Ordovician shales need be checked in the future. Anyway, the declining CIA values from the D. complanatus to P. pacificus biozones might suggest diminishing chemical weathering intensity (Figure 8). However, oxygen isotope evidence indicates that a mid-Boda cooling climate happened within the D. complanatus and D. complexus biozones and a late-Boda warming climate happened in the P. pacificus biozone (Figure 10) [93]. Moreover, the CIA trend is contrary to the observation that increasingly radiogenic ⁸⁷Sr/⁸⁶Sr occur within the P. pacificus biozone and indicate an augmented weathering flux (Figure 10). Therefore, this study believes that the decrease in CIA values and chemical weathering intensity in late Katian should not just be controlled by temperature regulation. Instead, tectonic and hydrologic regulations likely account for the changes in continental weathering. During the mid-Boda cooling interval, the chemical weathering and denudation rates were much lower due to the low temperatures and precipitation/runoff. In this interval, the low denudation rate would result in more clay formation and thick soils isolating the bedrock from weathering, which is the supply limited regime (high W/D and chemical weathering intensity; Figure 11).

A warming climate and orogenesis likely contributed to reducing chemical weathering intensity within the P. pacificus biozone. The warming climate is caused by global volcanism and the injection of CO₂ into the atmosphere (Figure 10; [94]), with the high temperatures in this interval leading to an increase in runoff of 6.8 to 2.3%/°C between high- and low-latitude rivers, respectively [95]. In this case, the rapid erosion of rocks and an exceptionally increased denudation rate would occur due to increased precipitation/runoff, as extensively reported in Early Silurian warming climates [17]. The orogeny in the late Katian has been reported globally, such as the Kwangsian Orogeny in south China, the Caledonian Orogen in Scotland, or the Taconic orogeny in the Appalachian margin of Laurentia, ensuring the continual exposure of fresh material [96,97,98], which tends to weather more rapidly. The intermediate denudation rate and enough fresh rocks would result in moderate erosion and intermediate weathering intensity (Figure 11).

Glaciation occurred within the M. extraordinarius and M. persculptus zones [93]. The weathering trends in Hirnantian Glaciation are subject to debate, with several perspectives reported. First, some studies propose that low temperatures and glacial ice cover could lower the chemical weathering rate, leading to a low chemical weathering intensity (low W/D ratio) [16]. Second, others suggest that glaciation enhances the denudation rate while reducing the chemical weathering intensity (low W/D ratio) [99]. However, the Yangtze Platform was located in subtropical regions during the Ordovician and hence was not subject to glacial weathering. Third, a cooling climate prolongs the continental residence time of water, allowing more clay formation, thereby reducing global weathering flux and promoting more incongruent weathering (intermediate W/D ratio) [27]. Given the high ⁸⁷Sr/⁸⁶Sr ratios and extremely low CIA values in Hirnantian Glaciation, this study suggests that the Hirnantian Glaciation was characterized by increased chemical weathering and denudation rates but low chemical weathering intensity (low W/D ratio). The amplified clastic fluxes from continents and increased continental material transported into the ocean are corroborated by lowεNd(t) values and high TiO₂/Al₂O₃, Zr/Al₂O₃ ratios [100]. Increased erosion during low sea level in glaciation, the glacial grinding of unweathered rocks, or a combination of them, caused the kinetically limited regime, which contributed to the increased denudation rate. Furthermore, the high seawater ⁸⁷Sr/⁸⁶Sr ratios and low sulfate–sulfur isotopes could support the increased chemical weathering. The high ⁸⁷Sr/⁸⁶Sr ratios might be driven by glacial grinding and abrasion that produced fine-grained Rb-rich glacial till with minerals such as biotite [101], as well as the preferential weathering of biotite with radiogenic ⁸⁷Sr/⁸⁶Sr in glaciation [101]. A large negative sulfur excursion (δ³⁴S in carbonate-associated sulfate) occurred throughout the Hirnantian Glaciation, and increasing continental weathering (especially sulfides) was required based on the geochemical box modeling [102]. Overall, the increased denudation rate exceeded the increased chemical weathering rate, leading to the lowest chemical weathering intensity in glaciation (Figure 11).

5.5. Implications for the Changes in Global Climates and LOMEs

The Late Ordovician and Early Silurian transition is characterized by an expansion of anoxia in the deeper oceans which resulted in the deposition of organic-rich shales which later became oil- and gas-source rocks [16,41,103]. Both the orbital forcing and paleoenvironmental changes have been explained for the deposition of those source rocks [9,104]. In the late Katian, a warming event (Boda event) was restricted to the upper part of the Ka4 time slice [105] or the late Katian P. pacificus biozone [63]. Previous findings imply that volcanism-induced paleoenvironmental repercussions in late Katian had long-term effects on the global system rather than short-term impacts on paleoenvironments and paleoclimates [106]. Therefore, volcanism likely contributed to the increase in pCO₂, negative δ¹³C excursion, and the late-Boda warming event [26]. In the P. pacificus interval, higher continental runoff is expected to have occurred, due to warming and a humid climate caused by intense volcanism. Furthermore, high runoff accelerates chemical weathering and denudation. The enhanced chemical weathering rate increased the consumption of atmospheric CO₂ and the nutrient influx into the ocean, leading to high primary productivity in surface waters [43,45]. The enhanced chemical weathering rate and weathering-induced high primary productivity contributed to the enrichment of organic matter, lower pCO_2, and lower global temperatures. Therefore, this study suggests that the weathering rate reaches maximum values for intermediate denudation rates and weathering intensity (Figure 11), which further contributes to CO₂ drawdown and cooling. Therefore, this study suggests that if intense volcanism can push the climate system out of balance (Boda warming), silicate weathering controlled by temperature and runoff would help to allow climatic recovery, even triggering glaciation [107]. We found that high chemical weathering and denudation rates, and the lowest chemical weathering intensity characterize the Hirnantian Glaciation. The high weathering rate seems to violate the fact that low temperatures and glacial cover characterized the Hirnantian. However, an increased chemical weathering rate in glaciations has also been proposed, with the factors being attributed to the enhanced erosion of silicate, and weathering of sulfides and carbonates [108]. In fact, the enhanced carbonate and sulfide weathering in the Hirnantian is supported by the high carbon isotopes and low sulfate-sulfur isotopes [102]. In this case, an increase in sulfide and carbonate weathering leads to a possible source of CO₂ followed by a longer sink of CO₂. This hypothesis is consistent with the observation of CO₂ release with the highest weathering and denudation rates (Figure 11). Overall, the highest denudation rate and lowest chemical weathering intensity in Hirnantian should result in the CO₂ release and termination of glaciation.

Traditional views suggested that the LOMEs were marked by two phases in the Late Ordovician (Figure 10). This study considered that the first phase corresponded to the start of glaciation and was within the boundary between Paraothograptus pacificus and Metabolograptus extraodinarius graptolite zones, and the second phase corresponded to the termination of glaciation and happened in the latest Hirnantian and earliest Rhuddanian periods (Figure 10; [4]). However, new studies found that the Late Ordovician extinctions happened within the middle–late Katian (named the Katian extinction) and also suggested a rapid global expansion of OMZs throughout the late Katian, based on trace elemental concentrations and Tl and U isotopes (Figure 10; [102]). This study suggests that weathering-induced high primary productivity and seawater anoxia likely played significant roles in controlling the LOMEs, given the coupling of high weathering rates and expanded anoxic seawater in the late Katian (Figure 10 and Figure 12). Previous studies suggested that oxygen level rather than temperature was the main factor controlling the Late Ordovician extinctions for as long as 35 million years [109]. The Late Ordovician crisis includes three rapid declines in species richness, with the first decline occurring in the late Katian, the second decline occurring in the earliest Hirnantian, and the third decline occurring in late Hirnantian (Figure 10; [110]). It is believed that weathering-induced seawater anoxia could explain the crisis in the late Katian, while the development of more intense euxinia and volcanism throughout the water column likely drove the second and third extinction in the Hirnantian (Figure 10; [14]).

6. Conclusions

The weathering trends during the Late Ordovician (late Katian-Hirnantian) have been evaluated by the chemical index of alteration (CIA) in shales and ⁸⁷Sr/⁸⁶Sr in carbonate leachates. The intervals within the D. complanatus and D. complexus biozones were marked by a low weathering rate (low ⁸⁷Sr/⁸⁶Sr ratios), and by high chemical weathering intensity (high CIA values). Conversely, the intervals within the P. pacificus biozone showed a high weathering rate (high ⁸⁷Sr/⁸⁶Sr ratios) and intermediate chemical weathering intensity (intermediate CIA values). This increase in weathering rate and reduction in chemical weathering intensity should be attributed to global orogenesis and enhanced precipitation/runoff under the warming climate. The Hirnantian Glaciation was characterized by high ⁸⁷Sr/⁸⁶Sr ratios and extremely low CIA values, which might be related to the exposure of continents during low sea level and the glacial grinding of unweathered rocks, with high ⁸⁷Sr/⁸⁶Sr ratios and low CIA values.

The changes in weathering had significant influences on the climatic changes and mass extinctions in the Late Ordovician. The intermediate weathering rate and intensity contributed to the CO₂ drawdown in the P. pacificus biozone, corresponding to the initiation of cooling and further glaciation. Meanwhile, the enhanced weathering-induced high primary productivity resulted in the expansion of anoxic seawater in the Katian, which further caused Katian extinction. In the Hirnantian Glaciation, the highest denudation rate and lowest chemical weathering intensity likely resulted in enhanced CO₂ release and contributed to the termination of glaciation.