Remagnetization of Carboniferous Limestone in the Zaduo Area, Eastern Qiangtang Terrane, and Its Tectonic Implications

Abstract

Robust paleomagnetic results through geological time are one of the keys to understand the drift history of the eastern Qiangtang terrane (EQT). Here, we presented comprehensive petrographic observations and rock magnetic and paleomagnetic analyses of the early Carboniferous Upper Zaduo (ZD) limestone Formation (C₁z₂) from the Sulucun (SLC) section in the Zaduo area, EQT, to investigate its magnetic originality and geological significance. A total of 12 sites (131 samples) were collected. Photomicrograph observations indicate that the limestone samples were characterized by widespread carbonate veinlets. Electron microprobe and energy dispersive spectrometry analyses confirm that authigenic magnetite formed after pyrite. Rock magnetic analyses reveal the dominant magnetic minerals of pyrite and magnetite, with ‘wasp-waisted’ hysteresis loops and close to the “remagnetization trend” hysteresis parameters. Based on both thermal and alternating field demagnetizations, the characteristic remanent magnetization directions for most samples were isolated: D_g = 6.3°, I_g = 50.1°, k_g = 54.9, α₉₅ = 6.2° in-situ, and D_s = 330.2°, I_s = 58.9°, k_s = 5.9, and α₉₅ = 20.5° after 2-step tilt correction. The κ (α₉₅) value decreases (increases) after tilt-correction, and the ChRM directions failed both the McFadden (1990), Watson and Enkin (1993) fold tests, indicating post-folding magnetizations. The 11 site-mean directions yield a mean in-situ paleopole of 84.4°N, 200.3°E, and A₉₅ = 6.8°, which is coincident with the post ∼53 Myr (especially around 40 Ma) paleopoles of the region. We therefore interpreted that these early Carboniferous limestone samples contain remagnetized magnetizations and that they were obtained after 53 Ma, most likely around 40 Ma, due to the far-field effect of the India–Eurasia collision.

Introduction

The Tibetan Plateau is composed of a series of continental terranes, such as from north to south, the Songpan-Ganzi terrane, the Qiangtang terrane (QT), the Lhasa terrane, and the Himalaya terrane (Figure 1A). These terranes drifted northward, and subsequent assemblages to Eurasia since the Late Paleozoic (Dewey et al., 1988; Yin and Harrison, 2000; Tapponnier et al., 2001; Metcalfe, 2013; Zhang et al., 2013; Zhu et al., 2013; Yan et al., 2016; Huang et al., 2018; Zhao et al., 2018; Song et al., 2020), which was accompanied by opening and closure of the Paleo-, Meso-, and Neo-Tethys oceans, have resulted in the formation and subsequent deformation of abundant marine carbonates preserved around the Tibetan Plateau (e.g., Yan et al., 2016; Guan et al., 2021). The QT, one of the major terranes in the central Tibetan Plateau (Figure 1A), is further divided into the eastern Qiangtang and western Qiangtang (alternatively called the “North Qiangtang terrane” and “southern Qiangtang terrane” in the literature, respectively) terrane by the so-called Longmo Co-Shuanghu suture zone (Figure 1A) (e.g., BGMRXAR, 1993; Yin and Harrison, 2000; Pan et al., 2004; Li et al., 2009; Zhu et al., 2013). Given that the EQT is immediately south of the Jinshajiang suture zone (Figure 1A), its drift history is important to understand the evolution and closure of the Paleo-Tethys Jinshajiang Ocean (Guan et al., 2021; Yu et al., 2022), as well as the tectonic history of the “Proto-Tibet”.

FIGURE 1

However, the Paleozoic affiliation and drift history of the EQT are still poorly constrained. For example, the presence of Carboniferous cool-water biotic assemblages and Carboniferous/Permian moraines of the Gondwana affinity suggests that the EQT was part of Gondwana during the period (Metcalfe, 2006; Metcalfe, 2013; Kent and Muttoni, 2020), while the presence of the Cathaysian fossils in the Permian limestone implies that the EQT had Cathaysian affinity and was away from Gondwana during the Carboniferous–Early Permian period (Liu and Sun, 2008; Zhang et al., 2013). In addition, some studies argued that the EQT originated from Laurasia, based on different geological and geophysical characteristics on both sides of the Bangong–Nujiang suture zone (Pan et al., 2004, 2012).

The large-scale latitudinal motions of terranes have made paleomagnetism one of the most important methods to decipher the opening and closure processes of the Tethys Ocean. The widely distributed limestone on the EQT has been one of the main targets for paleomagnetic studies to constrain its drift history. Many paleomagnetic studies have been carried out on limestone of the EQT, including rocks of the late Paleozoic (Cheng et al., 2012, Cheng et al., 2013; Yang et al., 2017) and the Mesozoic (Lin and Watts, 1988; Cheng et al., 2012; Ren et al., 2013; Yan et al., 2016; Ran et al., 2017; Cao et al., 2019; Zhou et al., 2019; Fu et al., 2021, 2022). Nevertheless, due to the complexity of limestone, some discordant results existed. For example, Cheng et al. (2012), Ren et al. (2013), and Yan et al. (2016) suggested primary magnetizations that yield similar paleomagnetic directions (paleolatitudes of ∼20–25°N) for the Middle-Upper Jurassic Yanshiping Group from the Yanshiping area, while Ran et al. (2017) argued for remagnetization directions; Cao et al. (2019) reported primary Middle Jurassic magnetizations of a different direction (paleolatitude of ∼35°N) from Shuanghu, whereas Fu et al. (2021) reported an Eocene remagnetization direction of the Middle-Upper Jurassic limestone from the Zaduo area. It is obvious that remagnetizations are common in the EQT that might have hindered our understanding on the drift history of the EQT and hence the evolution of the Paleo-Tethys Jinshajiang Ocean. Therefore, with more and more reported remagnetization results in the Qiangtang region, detailed analysis on the possibility of remagnetization is extremely important, especially for limestone.

In this study, we reported a detailed paleomagnetic study on the early Carboniferous limestone of the EQT, trying to provide a reliable Carboniferous paleomagnetic result. Unfortunately, based on detailed petrographic observations, rock magnetic experiments, and demagnetization analyses, this early Carboniferous limestone seems to have been remagnetized. We then discussed the possible acquisition mechanisms of remagnetization and its geological implication.

Geological Setting and Sampling

The EQT in the central Tibetan Plateau is separated from the Songpan-Ganzi terrane by the Jinshajiang suture zone to the north and from the western Qiangtang terrane by the Longmo Co-Shuanghu suture zone to the south. (Figure 1A) (e.g., BGMRXAR, 1993; Yin and Harrison, 2000; Zhu et al., 2013). Our study section is located ∼30 km south of Zaduo in the EQT (Figure 1A). In this area, the lithologic units mainly comprise the Lower Carboniferous Zaduo Group (limestone), the Upper Carboniferous Jiamainong Group (clastic rocks), the Triassic–Permian Gadikao Formation (tuff, rhyolite dacite, limestone, and siltstone), and the Middle-Upper Jurassic Yanshiping Group (cycles of clastic rock and limestone), above which overlie some Cenozoic deposits (QGSI, 2005a) (Figure 1B). The Zaduo Group can be further divided into two formations: the Lower Clastic and the Upper limestone ZD formations, with a conformable contact relationship (QGSI, 2014). Abundant fossils, including brachiopods (e.g., Gigantoproductus cf. giganteus (Sowerby)-Striatifera and Gigantoproductus edelburgensis-Sermiplanus Latssimlls), corals (Lithotrotion irregulare-Yuanophyllum sp., Yuanophyllum-Hxaphyllum, Lithotrotion irregulare Phillps, and Yuanophyllum sp.), and trilobites (Cummingella sp.), have been observed within the Zaduo Group, yielding the early Carboniferous (Tournaisian–Visean) age (QGSI, 2014). The Upper limestone ZD Formation is well exposed in the Zaduo area and extends more than 2,300 km² laterally, with a thickness up to 1,000 m. It formed in a shallow sea reef platform environment and mainly consists of bioclastic dolomitic limestone, cataclastic bioclastic dolomitic limestone, and bioclastic argillaceous limestone, intercalated with some siltstone and mudstone layers (QGSI, 2014).

We collected paleomagnetic samples from the Upper limestone ZD Formation at the Sulucun section (32.5°N, 95.1°E) along an anticline (part of the larger Siguo Syncline, Figures 2A,B,L) in the Zaduo area (QGSI, 2014). In this section, the formation mostly comprises bioclastic dolomitic/argillaceous limestone (Figure 2). The anticline has a NW-SE–trending axis plunging west (plunging direction = 275.8° and plunge angle = 20.3°), with its northeast limb dipping ∼49° and the southwest limb dipping ∼41–43° (Figure 2L). Given the presence of many faults and folds and some apparent deformation and absence of early Permian sediments in the study area (Figures 1B, Figure 2B), the exact time of folding is not precisely constrained. However, the Upper limestone ZD Formation is unconformably overlain by the folded Middle Jurassic Yanshiping Group. The fold axis of the Upper Carboniferous limestone is similar to that of the Middle Jurassic one (Figure 1B), likely indicating a post-Middle Jurassic folding. Our paleomagnetic sampling was carried out on both limbs along the Sulucun section, covering a range over 100 m. Ten to thirteen core samples were collected at each site by a portable gasoline-powered drill. The paleomagnetic samples were oriented with a magnetic compass and also a sun compass when weather allowed. The average difference between the magnetic and sun compass readings was 2.17° (n = 36, Supplementary Table S1); therefore, the magnetic compass can be used to constrain the direction of samples in this area. A total of 131 core samples were collected from 12 sampling sites (i.e., sites one to eight from the south limb and sites nine to twelve from the north limb) (Figure 2L). Most of the samples were collected from fresh rock away from apparent cracks and veins (Figure 2C).

FIGURE 2

Petrography

To better understand the texture, mineral composition, and potential alteration during burial, four representative samples (SLC1-3, 6-9, 8-3, and 9-3) from different parts of the limbs were microscopically observed by using a Zeiss Stemi 5,180 polarizing microscope and an electron microprobe analyzer (JXA-8230, JEOL, Japan) and were further analyzed by energy dispersive spectrometry analysis. These experiments were carried out in the State Key Laboratory of Tibetan Plateau Earth System, Resources and Environment, Beijing, China. In order to further determine the types of minerals, samples of SLC8-3 were analyzed by laser Raman spectroscopy on a Raman spectrometer (Renishaw inVia Qontor) in the Sichuan Keyuan Testing Center of Engineering Technology, Chengdu, China.

Photomicrographs in transmitted polarizing light reveal that two different microtextures are present in these four samples (Figures 2E–J). One is the bioclastic dolomitic limestone, with a granular structure, which consists mainly of bioclastic (account for 60%) and interstitial materials (40%). The bioclastics (e.g., ostracod, trilobite, foraminifera, brachiopod debris, algae, and moss) are almost completely altered by calcite, with a grain size of ∼0.10–1.80 mm (Figure 2E). The interstitial materials include calcite, dolomite, and terrigenous argillaceous with grain sizes of ∼0.01–0.1 mm. The other is the cataclastic bioclastic dolomitic limestone, which mainly consists of porphyroclasts (account for 45%) and interstitial materials (55%). The shape of the porphyroclast is irregular, and the grain size is mainly larger than 2.00 mm. Considering the shallow-water carbonate platform depositional environment, detrital inputs could therefore potentially contribute to the presence of porphyroclasts. The main component of interstitial materials is calcite, which has an irregular granular structure and a grain size of ∼0.1–1.60 mm. Two-stage carbonate veinlets are visible in two representative samples (SLC6-9 and SLC8-3, Figure 2F), indicating the presence of a diagenetic feature. Chalcedony, sphalerite, and a minor portion of oxidized pyrite were observed in SLC1-3 and SLC8-3 (Figures 2E,G,H). The presence of chalcedony and sphalerite may indicate a hydrothermal origin of the carbonate veinlets (Jiang et al., 2006; Yi et al., 2016), where the oxidization of pyrite may suggest a late diagenesis process. Partial dissolution and recrystallization are observed in all the four samples (e.g., Figures 2I,J). All these features may suggest the presence of hydrothermal events with the likelihood of remagnetization in the Upper limestone ZD Formation. In addition, the peak of sphalerite from sample SLC8-3 is very similar to that of standard sphalerite in the Raman spectrum (Figure 2K), which further indicates that the mineral is sphalerite.

Magnetite is the common mineral in the four representative samples. Two different morphologies of magnetic minerals are present (Figure 3). One phase is framboidal with individual framboid sizes ranging from only a few micrometers to >20 um (Figures 3H,M). Authigenic growth-zoning and rims are observed in some magnetite grains (Figure 3H). Another phase is subeuhedral to euhedral with small crystal sizes (<10 um); this population is usually distributed in the calcite matrix (Figure 3O). Meanwhile, pyrite (bright in backscattered electron images) is also present (Figures 3A–E,G–N,P). The pyrite grains are usually present as framboids (Figures 3G,I–K,N) or occur as the cores of subeuhedral to euhedral magnetite crystals (Figures 3A–E,L,M,P). Thus, the magnetite grains formed after pyrite are probably an oxidation product of pyrite during later diagenetic events (Suk et al., 1990). In addition, rutile grains also presented in two samples (SLC6-9 and SLC9-3, Figures 3F,N). They are usually the diagenetic alteration product of detrital titanomagnetite with iron completely leached from the particles (Huang et al., 2017). Therefore, the magnetite grains are authigenic that likely formed as an oxidation product of pyrite or other iron sulfides and are hence apparently secondary.

FIGURE 3

Energy dispersive spectrometry analysis indicates the main mineral composition of iron, oxygen, sulfur, and titanium (Figure 4). This further supports the aforementioned observations in the presence of magnetite, pyrite, and rutile.

FIGURE 4

Paleomagnetic Measurements and Results

Measurement Methods

To characterize the magnetic properties of the samples, a series of rock magnetic experiments were conducted. High-temperature magnetic susceptibility measurements were performed on six representative samples by using a MFK1-FA Kappabridge with a CS-4 high-temperature furnace. Seven cycles of successive heating to 250, 350, 400, 450, 550, 620, and 700°C with intervening cooling to room temperature were carried out on each sample in an argon atmosphere. Hysteresis loop, IRM acquisition, and reverse field demagnetization curve measurements of 69 samples (five to seven representative samples/site) were measured on a Lakeshore 8600 vibrating sample magnetometer at room temperature. After using an ASC IM-10-30 pulse magnetizer with a maximum applied field of ∼100 mT, IRM acquisition curves for eight representative samples were measured by using a Minispin magnetometer. Most of the samples from the Upper limestone ZD Formation were first thermally demagnetized and then followed by alternating field demagnetization to isolate the ChRM directions. Some samples were only subjected to progressive thermal or alternating field demagnetizations. Additionally, both demagnetizations were carried out on pared sister specimens (Supplementary Table S2). Thermal demagnetization was carried out in an ASC TD-48 furnace (with a residual field <10nT). Alternating field demagnetization was performed up to 140mT by a degausser attached to the 2G Enterprises Model 755 cryogenic magnetometer (RAPID system). Most of the samples were progressively demagnetized in 24 steps (i.e., NRM, 80, 120, 150, 180, 200°C, 2.5mT, 5mT, 7.5mT, 10mT, 12mT, 15mT, 20mT, 25mT, 30mT, 40mT, 50mT, 60mT, 70mT, 80mT, 90mT, 100mT, 120mT, and 140mT), and the sister specimens were either progressively thermal demagnetized in 24 steps (NRM, 130, 170, 210, 250, 290, 320, 340, 360, 380, 400, 410, 415, 420, 425, 430, 435, 440, 445, 450, 455, 460, 465, and 470°C) or alternating field demagnetized in 19 steps (NRM, 2.5mT, 5mT, 7.5mT, 10mT, 12mT, 15mT, 20mT, 25mT, 30mT, 40mT, 50mT, 60mT, 70mT, 80mT, 90mT, 100mT, 120mT, and 140mT). All measurements were carried out in a magnetically shielded room that has an average field of less than 170 nT in the paleomagnetic laboratory in the State Key Laboratory of Tibetan Plateau Earth System, Resources and Environment, Beijing, China.

Rock Magnetism

High-Temperature Susceptibility

Stepwise thermomagnetic runs of magnetic susceptibility were carried out (Figure 5). All of the 250°C and some of the 400°C (Figure 5A,D–F) heating–cooling cycles are reversible, which may suggest no mineral transformation. Some curves are characterized by an increase in susceptibility after 420–550°C (Figure 5D–F), implying formation of new ferrimagnetic minerals during heating (e.g., pyrite, Deng et al., 2001; Zan et al., 2017). Some samples show a decrease in the susceptibility at ∼580°C and are followed by a slow decrease up to ∼680°C (Figure 5C,D), suggesting the presence of both magnetite and hematite. Meanwhile, most of the heating–cooling cycles are quasi-reversible during a higher temperature interval (i.e., 700°C) (Figure 5B, D–F), with a distinct hump at ∼450–550°C, which is likely the Hopkinson peak (Figure5D–F).

FIGURE 5

Hysteresis Loop, IRM Acquisition, and IRM Component Analyses

Hysteresis loops show a typical ‘wasp-waisted’ feature for most samples (Figures 6A,C–F), which is also indicated by quantitative analysis of that of Fabian (2003) that E_hys is larger than 4M_sB_c. This feature indicates the existence of multiple magnetic components with distinct coercivities, which may correspond to the mixtures of different coercivity magnetic minerals or different size fractions of a single mineral (Tauxe et al., 1996; Jackson and Swanson-Hysell, 2012). Most of the specimens (except SLC9-7, remanent coercivities: 83mT, Figure 6F) were saturated below 0.5 T, together with low coercive forces (B_c, 4.02–7.87mT) and remanent coercivities (B_cr, 34–46mT), indicating that the dominant magnetic carrier is probably magnetite (some samples also with hematite). The IRM acquisition curves reveal that their remanences are ∼64–72% saturated by 0.1T and almost fully saturated (∼92–96%) at 0.3T (Figures 6G–I), implying the major remanent carrier of the “soft” magnetic phase, such as magnetite. This is further supported by the relatively narrow hysteresis loops (Figures 6A–F). Decomposition of the IRM acquisition data of three representative samples reveals a similar four-component model (Maxbauer et al., 2016) fit (Figures 6J–L). The low-coercivity component (components 1 and 2) has B_h (the mean coercivity of an individual grain population) values of ∼0.69–1.34 log₁₀ units (∼5–22mT) and DP (dispersion parameter) values of ∼0.2–0.3, ∼6–22% contribution. Component 3 has B_h values of ∼1.78–1.85 log₁₀ units (∼60–70mT) and DP values of ∼0.3, >62% contribution, indicating the dominance of magnetite. Component 4 has B_1/2 values of ∼2.48–2.58 log₁₀ units (∼300–380mT) and DP values of ∼0.2, ∼2–4% contribution, implying the presence of very fine-grained magnetite that is close to the threshold size of SP, or residual iron sulfide after oxidizing to authigenic magnetite (Figures 4J–L).

FIGURE 6

Day Plot

Although non-definitive, the Day plot provides an indicative means of discriminating secondary from primary magnetite in limestone (Meng et al., 2020). In Figure 7, the hysteresis parameters of 69 samples are displayed in the Dunlop (2002)’s mixture zone of the pseudosingle domain (PSD), superparamagnetic (SP) to single domain (SD) magnetite. They differ from the range of primary magnetization carbonates but are rather close to the range of widespread remagnetized carbonates (Jackson and Swanson-Hysell, 2012; Fu et al., 2021), where the B_cr/B_c ratios range from 2 to 18 and M_rs/M_s ratios range from 0.08 to 0.2 (Supplementary Table S3, Figure 7). The low B_c and B_cr values indicate that these ‘wasp-waisted’ loops likely represent a mixture result of SD/PSD and SP magnetite, or in some cases possibly also a mixture of both hard and soft magnetic minerals. Although, most natural sedimentary samples that fall within the PSD domain make the interpretation more complicated (e.g., Qin et al., 2008; Cao et al., 2019; Fu et al., 2021; Guan et al., 2021), some still pertain to the Day plot that remagnetization may be diagnosed (Roberts et al., 2018).

FIGURE 7

Paleomagnetic Directions

Principal component analysis involving at least four successive steps was employed to determine the magnetization directions by PaleoMag software (v. 3.1d40) of Jones (2002). Specimens with maximum angular deviation (MAD) > 15° and sites with sample number <5 and/or α₉₅ > 16° were rejected for further analysis. Site-mean directions were calculated by standard Fisher statistics (Fisher, 1953). Most of the specimens exhibited a single component that linearly decays to the origin. This component can generally be isolated below 470°C or 140mT. In addition, both the thermal and alternating field demagnetization results of a same sample share a similar direction (Figure 8, Supplementary Table S2). Hence, a Fisherian site mean is calculated by both the thermal and alternating field demagnetization specimens.

FIGURE 8

A total of 104 out of 131 samples (11 out of the 12 sites) have stable ChRM directions. The obtained overall mean direction is D_g = 6.3°, I_g = 50.1°, k_g = 54.9, α₉₅ = 6.2° in-situ, and D_s = 330.2°, I_s = 58.9°, k_s = 5.9, and α₉₅ = 20.5° after the 2-step tilt correction (Stewart and Jackson, 1995) (Figures 9A,B). The κ (α₉₅) value decreases (increases) after the tilt correction. These ChRM directions fail both the Watson and Enkin (1993) progressive unfolding test (unfolding k_max = 66.54 at 12.4%, with 95% uncertainties ranging from 5.6 to 18.8%) (Figure 9C) and McFadden (1990) fold test (ξ1 = 3.3 before and ξ1 = 7.8 after tilt correction, with critical values of ξ = 3.87 at 95% and ξ = 5.38 at 99% confidence levels), indicating post-folding magnetizations. The 11 in-situ site-means, when converted to virtual geomagnetic poles, yield a mean in situ paleopole of 84.4N, 200.3°E, and A₉₅ = 6.8°, corresponding to a paleolatitude of 30.9°N for the sampling area (Table 1, 2).

FIGURE 9

Discussion

Evidence for Remagnetizations

In the study area, as mentioned earlier, carbonate veinlets are widespread in the Upper limestone ZD Formation, such as a few to a dozen millimeters wide on the surface of some samples and a few to tens of microns wide under a microscope (Figures 2D–F,H). These carbonate veinlets are mainly composed of calcite, commonly believed as an indicator of diagenetic fluid migration (Gustavson et al., 1994; Phillip, 2008; Bons et al., 2012; Gale et al., 2014), yielding likelihood of remagnetization in the strata. Energy dispersive spectrometry and electron microprobe observations of the representative samples show that the magnetite grains have authigenic growth-zoning and rims, where the pyrite grains usually occur as framboids or the cores of magnetite crystals, with the presence of some rutile grains (Figures 3, 4). It suggested that the magnetite grains are authigenic that likely formed as an oxidation product of pyrite or other iron sulfides.

Authigenic magnetite (dominantly SP and SSD) is rather commonly used to be indicative of the occurrence of remagnetization (McCabe and Channell, 1994; Jackson and Swanson-Hysell, 2012). Our rock magnetic results show that the dominant magnetic carriers in the Upper limestone ZD Formation are pyrite and magnetite (some samples also with hematite) (Figures 5, 6). The hysteresis loops indicate the ‘wasp-waisted’ feature in most samples (Figures 6A,C–F). It is worth noting that the alternating field demagnetization and some rock magnetism results (Figures 5, 6) indicate the presence of high coercivity magnetic minerals, e.g., hematite. However, they have a relatively low content, as indicated by low unblocking temperatures in thermal demagnetization processes and no presence of hematite in all representative samples in petrographic observations (Figures 3, 8). Anyway, the presence of ‘wasp-waisted’ loops is most likely due to the mixture of SP and SSD grains of magnetite (e.g., Dekkers and Pietersen, 1991; Tauxe et al., 1996; Gong et al., 2009), or only a few cases with the mixture of hard and soft magnetic minerals (e.g., hematite and magnetite), or a combination of both situations in few cases. In addition, the unblocking temperature depends on many factors (e.g., the type, volume, shape, and element content of magnetic minerals), and the low unblocking temperature (400–500°C) may indicate the presence of fine-grained magnetite or titanomagnetite (O'Reilly, 1984; Liu et al., 2007) (Figure 8, Supplementary Table S2). Meanwhile, the coercivity (B_cr/B_c) and remanence (M_rs/M_s) ratios are rather close to the “remagnetization trend” on the Day plot; a zone was previously interpreted to be the characteristic of chemical remagnetization, which is distinct from the primary limestone magnetization region (e.g., Jackson 1990; McCabe and Channel, 1994; Jackson and Swanson-Hysell, 2012; Fu et al., 2021) (Figure 7). Furthermore, our result shows a similar trendline with the empirically derived equation M_rs/M_s = 0.89(B_cr/B_c)^−0.6 (Jackson, 1990; Jackson et al., 1993) of the uncommon magnetic properties of remagnetized carbonates, except for the lower M_rs/M_s value, which can be attributed to the partial oxidation of magnetite, particle shape, or uncertain mixtures of magnetic minerals (Roberts et al., 2018). The result fits well when changing the equation to M_rs/M_s = 0.89(B_cr/B_c)^−l or M_rs/M_s = 0.5(B_cr/B_c)^−0.6, where the best fit equation of our data is M_rs/M_s = 0.37(B_cr/B_c)^−0.5 (Figure 10).

FIGURE 10

In addition, the obtained ChRMs of the Upper limestone ZD Formation yield a negative fold test, suggesting post-folding magnetizations, which likely occurred after the Middle Jurassic. Meanwhile, the 11 2-step tilt-corrected site-means provide a mean paleopole of 65°N, 30.2°E, and A₉₅ = 26.4° that corresponds to a paleolatitude of 39.7°N for the sampling area (Table 2). This is discordant to the available geological and paleomagnetic results that the EQT was located in the southern hemisphere during the late Paleozoic (Zhao et al., 1996; Li et al., 1999; Song et al., 2017; Huang et al., 2018; Zhao et al., 2018; Guan et al., 2021), rather than at such middle latitude in the northern hemisphere.

In summary, the aforementioned petrographic observations and rock magnetic and paleomagnetic measurements suggest that these early Carboniferous limestone samples most likely contain remagnetized magnetizations. Authigenic SP and SSD magnetite and/or a mixture of hematite and magnetite are responsible for the secondary magnetizations.

Timing and Mechanism of Remagnetization

In principle, remagnetization can occur at any time during the geological history, yet it is generally related to major tectonic events (e.g., orogeny and metamorphism). There are four major tectonic events reported in the Tibetan Plateau region after the early Carboniferous, such as the collisions of the EQT with the Tarim/Songpan-Ganzi/Yidun terranes, the western Qiangtang terrane with the EQT, the Lhasa with the western Qiangtang terrane, and India with Eurasia plates. Our negative fold test of the obtained ChRMs indicates post-folding magnetizations, which extensively occurred after the Middle Jurassic. Thus, the remanence was likely obtained after the Middle Jurassic. However, the time of folding is rather extensive without robust geological evidence, and the post-Middle Jurassic is a relatively long period; further analyses are essential to constrain the time of remagnetization.

The shortest distance from a remagnetization paleopole to the reliable reference poles is commonly used to estimate the time of the remagnetization event in paleomagnetism (Van der Voo and Torsvik, 2012). We collected all the available post Carboniferous paleopoles of the EQT and filled with the quality criteria (Van der Voo, 1990; Meert et al., 2020), as shown in Figure 11 and Table 2. Paleopoles are relatively concordant at each period, such as during the Permian (Song et al., 2017; Ma et al., 2019; Guan et al., 2021), the Triassic (Song et al., 2015, 2020; Yu et al., 2022), the Jurassic (Cheng et al., 2012; Yan et al., 2016), and the Cretaceous (Huang et al., 1992; Tong et al., 2015; Meng et al., 2018), except that the Cenozoic poles are rather scattered (Figure 11), which are likely due to widespread local rotations during the Lhasa–Qiangtang terrane and India–Eurasia collisions (Tong et al., 2015; Chen et al., 2017). Our obtained in-situ paleopole of 84.4°N, 200.3°E with A₉₅ = 6.8° is far away from these known Paleozoic and Mesozoic poles but rather close to ∼53–38.6 Ma poles (especially around 40 Ma, Zhang et al., 2020; Lippert et al., 2011) and the present day pole (Figure 11; Table 2), while given that the remagnetization time of the Middle-Upper Jurassic limestone was proposed to be Eocene in the adjacent area of only ∼8 km away (Figure 1B) (Fu et al., 2021). Hence, both of the two strata likely had experienced similar Eocene remagnetizations. We therefore interpreted the episode of remagnetization of the early Carboniferous limestone (the Upper limestone ZD Formation) to be after 53 Ma, most likely around 40 Ma, during the early stage of the India–Eurasia collision, although any post 53-Myr remagnetization still cannot be ruled out.

FIGURE 11

A further question is how were these limestones remagnetized? The most common mechanisms to explain remagnetization include the thermoviscous resetting of existing magnetic minerals (Kent, 1985) and chemical remanent magnetization through magnetic mineral growth associated with orogenic fluids (e.g., Jackson, 1990; Elmore et al., 2012; Huang et al., 2015). Given that only the Late Triassic and Early Cretaceous intrusive rocks are found near the section (Figure 1B), they are deemed a less likely mechanism for the remagnetization as 1) the obtained pole is most likely of the Early Cenozoic one; 2) The nearby Permo-Triassic volcanic rocks and Early Cretaceous intrusive rocks both contain primary magnetizations (Guan et al., 2021; Fu et al., 2022). In general, limestone perhaps has not been heated for a sufficient amount of time at such a temperature to make the thermal resetting feasible by the intrusive rocks (Dekkers, 2012); and 3) The nearby Middle Jurassic limestone was remagnetized during the Eocene (Fu et al., 2021).

The Upper limestone ZD Formation was deposited in a littoral and shallow sea carbonate platform environment (QGSI, 2014). The appearance of massive bioclastic limestone indicates a warm climate at that time, which was conducive to biological growth (QGSI, 2014). Previous studies suggested that when organic carbon fluxes were high, oxygen would be used up and form an anoxic sulphidic diagenetic environment, where paramagnetic pyrite would have to be replaced with detrital magnetite and hematite (Froelich et al., 1979; Roberts, 2015; Huang et al., 2019; Fu et al., 2021). Thus, the Upper limestone ZD Formation tended to contain plenty of iron sulphides during its burial and diagenesis process. Late, the Qamdo region has been uplifted to a fairly high elevation during the early stage of the India–Eurasia collision around the Paleogene (Xu et al., 2013; Tang et al., 2017; Xiong et al., 2020). The environment might have turned to suboxic and/or oxic during the period, resulting in iron sulphide oxidation to authigenic magnetite and the acquisition of chemical remanent magnetizations (Brothers et al., 1996). Meanwhile, the process might have caused migration of orogenic fluids (e.g., hydrothermal, Figure 2), resulting in the occurrence of widespread carbonate veinlets in the Upper limestone ZD Formation and the nearby Middle-Upper Jurassic limestone (Fu et al., 2021), leading to the chemical remanent magnetizations of the previous Upper Carboniferous and Middle-Upper Jurassic limestones.

Therefore, the obtained paleomagnetic direction of the Upper limestone ZD Formation is a remagnetized direction. Given both the well-observed primary and secondary directions in limestone samples elsewhere around the Tibetan Plateau (e.g., Lin and Watts, 1988; Yan et al., 2016; Huang et al., 2017; Ran et al., 2017; Zhao et al., 2021), paleomagnetic directions recorded in the limestone are extremely complicated around the region. It is necessary to carry out detailed rock magnetic and petrographic analyses to obtain robust paleopoles in the Tibetan Plateau.

Conclusion

Petrographic observations of the Upper limestone ZD Formation demonstrate the magnetic minerals of pyrite and authigenic magnetite. The presence of chalcedony and sphalerite suggests that these representative samples have more likely been affected by hydrothermal activities. High-temperature susceptibility and hysteresis loop analyses reveal the dominant magnetic minerals of pyrite and fine-grained magnetite; the coercivity (B_cr/B_c) and remanence (M_rs/M_s) ratios are rather close to the ‘remagnetization trend’. Paleomagnetic demagnetization analyses obtained 11 sites of 104 ChRM directions, which failed the fold test, indicating post-folding magnetizations. These ChRMs yield an in-situ paleopole of 84.4N, 200.3°E, and A₉₅ = 6.8°, which is coincident with the post ∼53 Myr (especially around 40 Ma) paleopoles of the region. We ascertained that these early Carboniferous limestones contain remagnetized magnetizations that were obtained after 53 Ma, most likely around 40 Ma, due to the far-field effect of the India–Eurasia collision.

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