The quantitative analysis of in situ major elements of minerals was completed by using an electron probe microanalyzer at Wuhan SampleSolution Analytical Technology Co., Ltd., Wuhan, China. Complete the analysis using JXA-8230 of JEOL. The voltage and current analyzed are 15 KV and 10 nA, respectively, with a beam diameter set to 3 μm for feldspar and mica and 1 μm for epidote. The calibration standard samples for the content of major elements use 53 kinds of mineral standard samples, 44 kinds of elemental standard samples, and 15 kinds of rare earth element standard samples provided by SPI Company. The data correction method adopts the ZAF correction method of JEOL. The results are listed in Supplementary Table S1.

The sample pretreatment of whole rock for major element analysis was made by the melting method. The flux is a mixture of lithium tetraborate, lithium metaborate, and lithium fluoride (45:10:5). Ammonium nitrate and lithium bromide were used as oxidants and release agents, respectively. The melting temperature was 1,050°C and the melting time was 15 min. Zsx Primus II wavelength dispersive X-ray fluorescence spectrometer (XRF) produced by RIGAKU, Japan was used for the analysis of major elements in the whole rock. The data were corrected by the theoretical α coefficient method. The relative standard deviation (RSD) is less than 2%. The whole-rock geochemical data are presented in Table 1.

U–Pb dating and trace element analysis of zircon were simultaneously conducted by LA-ICP-MS at the Wuhan SampleSolution Analytical Technology Co., Ltd., Wuhan, China. An Agilent 7900 ICP-MS instrument was used to acquire ion-signal intensities. The spot size and frequency of the laser were set to 24 µm and 5Hz, respectively. Zircon Tanz and glass NIST610 served as external standards for U–Pb dating and trace element calibration (Hu et al., 2021), respectively. Each analysis incorporated a background acquisition of approximately 20–30 s, followed by 50 s of data acquisition from the sample. An Excel-based software, ICPMSDataCal, was used to perform off-line selection and integration of background and analyzed signals, time-drift correction, and quantitative calibration for trace element analysis and U–Pb dating (Liu et al., 2008). Concordia diagrams and weighted mean calculations were conducted using Isoplot/Ex_ver3 (Ludwig, 2003). The results are listed in Supplementary Table S2.

The investigated suite comprises biotite-plagioclase gneisses and biotite schist sampled from the Sijiaying iron ore district, Xiangtang Town, Luanzhou City. These supracrustal units are stratigraphically interlayered with BIFs, which exhibit characteristic rhythmic alternations of siliceous and iron-rich laminae (Figure 2a). Within the biotite schist, a penetrative tectonic foliation is defined by preferentially oriented biotite aggregates, displaying continuous planar alignment wherein quartz, plagioclase, and K-feldspar form elongate grains parallel to the foliation plane (Figure 2b). Sample N16-1 (biotite-plagioclase gneiss) manifests typical gneissic banding with irregular lenticular segregations of biotite-dominated mafic layers contrasting against felsic mineral domains (Figure 2c). While sharing similar structural characteristics, sample N16-6 exhibits coarser-grained gneissose banding relative to N16-1 (Figure 2d).

The main mineral components in biotite schist (N12-6) include biotite (65 vol%), K-feldspar (14%), plagioclase (3%), epidote (8%), and quartz (10%), with minor accessory minerals such as zircon and apatite (Figure 2e). The K-feldspar grains are uniformly sized (0.1–0.2 mm), appearing subhedral to anhedral. Chemically, the K-feldspars are dominated by microcline (Figure 3a). Biotite can be divided into two types based on grain size. The coarse-grained biotite is of 0.4–0.7 mm across, and the fine-grained one is 0.1–0.3 mm across (Figure 2h). Chemical composition analysis reveals that the X_Mg [= Mg/(Fe²⁺ + Mg)] of biotite shows limited variation (0.39–0.41). The X_Ti content in fine-grained biotite (0.03–0.05 a.p.f.u.) is slightly lower than those in coarse-grained biotite (0.04–0.07 a.p.f.u.) (Figure 3b). Both types are classified as ferroan biotite based on their compositional plots (Figure 3c). Epidote occurs either as cross-cutting veins through biotite or as disseminated grains within the matrix.

The biotite-plagioclase gneiss (N16-1) exhibits a grayish-black color with gneissic structure characterized by oriented biotite flakes intergrown with granular minerals (Figure 2f). The mineral assemblage comprises biotite (25 vol%), K-feldspar (2%), plagioclase (34%), epidote (3%), and quartz (36%), along with accessory muscovite, sphene, zircon and calcite (Figure 2i). Biotite grains show uniform size. Microprobe traverses of biotite exhibit zoning profiles, with the X_Ti contents increasing from core to rim (0.085–0.102), while X_Mg values (0.52–0.53) in biotite is homogeneous. According to compositional discrimination diagrams, they were classified as magnesian biotite (Figure 3c). K-feldspar occurs as xenomorphic to subhedral microcline, while plagioclase is predominantly oligoclase with X_An = 0.20–0.21. Some albite grains occur as inclusions in biotite. Epidote occurs as fine-grained aggregates in the matrix.

The mineral assemblage in sample N16-6 consists mainly of biotite (18 vol%), K-feldspar (9%), plagioclase (43%), muscovite (1%), and quartz (29%), along with minor epidote, albite and zircon (Figure 2g). The biotite can be divided into coarse-grained biotite (0.3–0.4 mm) and fine-grained biotite (<0.2 mm) based on grain size (Figure 2j). The X_Ti contents in coarse-grained biotite generally decrease (0.104–0.088) from core to rim, while X_Ti contents in fine-grained biotite is relatively homogeneous (0.099–0.090). Both types exhibit X_Mg values of 0.49–0.51, indicating magnesian biotite based on compositional classification diagrams (Figure 3c). The K-feldspar has a grain size of 0.2–0.3 mm, all being microcline. Plagioclase grains are across of 0.2–0.5 mm, being oligoclase with X_An = 0.17–0.23. Minor fine-grained epidote and albite occurs as inclusions within biotite grains, interpreted to have formed in a later stage than the matrix biotite.

The representative supracrustal rock samples collected from Sijiaying area include biotite schist and biotite-plagioclase gneisses. The biotite schist (N12-6) has a SiO₂ content of 43.43 wt%, which is lower than that of the biotite-plagioclase gneisses. It exhibits higher Al₂O₃ (16.12 wt.%) and Fe₂O₃ ^T (19.76 wt.%), with MgO = 6.22 wt.%, CaO = 3.49 wt.%, X_Mg = 0.26, and A/CNK = 1.45 (Table 1). Combined with protolith reconstruction diagrams, these data indicate that the protolith of this schist is shale formed in an Fe-Al-rich sedimentary environment (Figure 3d). The biotite-plagioclase gneisses have SiO₂ contents of 65.89–67.45 wt.%, Al₂O₃ = 12.73–14.89 wt.%, Fe₂O₃ ^T = 4.34–5.95 wt.%, MgO = 2.19–2.91 wt.%, CaO = 1.39–2.56 wt.%, X_Mg = 0.35–0.40, and A/CNK = 1.48–1.53. The protoliths of these gneisses are classified as greywacke on the discrimination diagram (Figure 3d).

The zircon grains in sample N16-1 are predominantly subhedral to anhedral, displaying prismatic or elliptical shapes (Figure 4a), with aspect ratios of 1:1–1:3 and long-axis lengths of 50–130 μm. Most zircons exhibit core-rim structures, where the cores show distinct or blurred oscillatory zoning, suggesting their magmatic orogin (Wu and Zheng, 2004). While a few displays planar zoning or without zoning. The rims are extremely narrow and dark gray. A total of 70 analytical spots on 62 zircon grains were analyzed for trace element compositions and U–Pb dating. The Th/U ratios range from 0.05 to 1.23. In the ²⁰⁷Pb/²³⁵U-²⁰⁶Pb/²³⁸U concordia diagram (Figure 4c), most zircons deviate from the concordia line due to varying degrees of Pb loss. The ²⁰⁷Pb/²⁰⁶Pb apparent ages range from 1731 ± 55 Ma to 2,898 ± 59 Ma, with a peak age of 2,552 Ma. Analyses plotting on or near the concordia line yield a weighted mean ²⁰⁷Pb/²⁰⁶Pb age of 2,547 ± 14 Ma (MSWD = 0.32, n = 29).

The zircon grains in sample N16-6 are predominantly subhedral to anhedral, displaying prismatic or elliptical shapes (Figure 4b), with aspect ratios of 1:1–1:2 and sizes of 60–120 μm. Most zircons exhibit core-rim structures, where the cores show distinct or blurred oscillatory zoning, while a few displays weak planar zoning with bright luminescence. The rims are extremely narrow and light gray. A total of 73 analytical spots on 64 zircon grains were analyzed for trace element compositions and U–Pb dating. The Th/U ratios range from 0.18 to 1.61. In the ²⁰⁷Pb/²³⁵U-²⁰⁶Pb/²³⁸U concordia diagram (Figure 4d), most zircons deviate from the concordia line due to varying degrees of Pb loss. The ²⁰⁷Pb/²⁰⁶Pb apparent ages range from 2,229 ± 67 Ma to 2,727 ± 37 Ma, with a peak age of 2,556 Ma. Analyses plotting on or near the concordia line yield a weighted mean ²⁰⁷Pb/²⁰⁶Pb age of 2,555 ± 14 Ma (MSWD = 0.30, n = 41). Chondrite-normalized REE patterns show that zircon grains from samples N16-1 and N16-6 predominantly display positive Ce anomalies, (Lu/Gd)_N values of 2.22–89.86, and HREE contents higher than LREE, resulting in steeply left-inclined patterns (Figures 4e,f).

Phase quilibria are modelled for samples N16-1 and N16-6 in the system MnNCKFMASHTO (MnO-Na₂O-CaO-K₂O-FeO-MgO-Al₂O₃-SiO₂-H₂O-TiO₂-O). Calculations are performed using GeoPs software (Xiang and Connolly, 2021) incorporating the ds62 thermodynamic dataset (Holland and Powell, 2011). Mineral and activity-composition models used in the calculations are plagioclase and K-feldspar (Holland and Powell, 2003), biotite and muscovite (White et al., 2014), epidote (Holland and Powell, 1998), ilmenite and magnetite (White, 2000), garnet, chlorite, cordierite and melt (White et al., 2014). Sillimanite, andalusite, rutile, titanite, and quartz were pure end-member components. P–T pseudosection for each sample was calculated using an effective bulk-rock composition which was generated according to mass balance constraints by integrating mineral compositions and modal abundance data of the phases present (Carson et al., 1999). The effective bulk-rock compositions used for phase diagram calculations are listed in Table 1.

Based on the petrographic characteristics and phase equilibria modelling, the metamorphic evolution of the supracrustal rocks from Sijiaying area, eastern Hebei, includes three stages: (i) pre-peak prograde to peak stage; (ii) peak stage; and (iii) post-peak decompression and cooling stage. For sample N16-6, a clockwise P–T path involving peak condition and post peak decompression and cooling process was well defined. The peak P–T condition was constrained to be at ∼7.0 kbar/630°C based on the minimum X_An in plagioclase and maximum X_Ti isopleths in biotite, followed by decompression and cooling with metamorphic reaction of pl + mus = bt + kfs + qz + H₂O. The pre-peak prograde process can be inferred from the inclusion assemblages of epidote, sphene and albite within biotite and K-feldspar, which was dominated by metamorphic reaction of mus + ab + ep + sph + qz = kfs + bt + pl + H₂O. For sample N16-1, the pre-peak prograde to peak stage is inferred on the basis of biotite zoning and inclusions in matrix biotite and feldspar, showing a P–T segment with increase in both pressure and temperature, dominated by metamorphic reaction of ab + sph + ep + mus + qz = Pl + bt + kfs + H₂O. The peak condition is defined by the minimum X_An in plagioclase and maximum X_Ti isopleths in biotite. The post-peak stage, including decompression process, in sample N16-1 is inferred to be similar as those in sample N16-6. Phase equilibria modelling for sample N12-6 (biotite schist) defines broadly constrained P–T conditions, precluding precise determination of peak metamorphism. However, the X_Ti in biotite (0.03–0.07 a.p.f.u.) from biotite schist is markedly lower than that of the two biotite-plagioclase gneiss samples, suggesting greenschist-facies metamorphism indicative of shallower crustal depths relative to the gneissic units.

Previous geochronological studies on the eastern Hebei terrane demonstrate that the protolith of metamorphic supracrustal rocks predominantly formed between 2.61 and 2.50 Ga (Guo et al., 2013; Wan et al., 2015; Sun et al., 2016; Fu et al., 2016; Lu et al., 2017; Duan et al., 2017; Liu and Wei, 2020), with magmatic activity concentrated at 2.56–2.48 Ga (Geng et al., 2006; Nutman et al., 2011; Bai et al., 2016; Yang and Wei, 2017b; Li et al., 2019; Duan et al., 2022). The eastern Hebei terrane underwent two phases of metamorphism during the Neoarchean and Paleoproterozoic, with metamorphic ages primarily concentrated at 2.53–2.47 Ga and 1.85–1.77 Ga (Geng et al., 2006; Nutman et al., 2011; Yang and Wei, 2017b; Duan et al., 2019; Liu et al., 2020; Li et al., 2024). The biotite-plagioclase gneisses of the Luanxian Group in this study contain magmatic zircon grains with significantly old ²⁰⁷Pb/²⁰⁶Pb apparent ages of 2.90–2.63 Ga, indicating the presence of ancient rocks or zircons in the source region. Some ancient age records have also been reported from Labashan area (>3.4 Ga) in Lulong County (Chu et al., 2016; Wan et al., 2021; Zhao et al., 2023; Dong et al., 2024), Zhuzhangzi area (>2.90 Ga) in Qinglong County (Sun et al., 2010; Guo et al., 2015) and Huangbaiyu area in Qian’an (>3.8 Ga) (Liu et al., 1992; Wu et al., 2005; Wilde et al., 2008), indicating prolonged and widespread Archean magmatism. The youngest concordant ²⁰⁷Pb/²⁰⁶Pb age group together with weight mean ages of 2,547 ± 14 Ma and 2,555 ± 14 Ma from biotite-plagioclase gneiss provides a maximum age for the deposition of the Luanxian supracrustal rocks in this study, which is consistent with the depositional age of the Sijiaying BIFs (Cui et al., 2014; Han et al., 2014; Fu et al., 2021; Gao et al., 2023). Metamorphic volcanic rocks and BIFs from the Luanxian Group record ∼2.50 Ga metamorphic ages (Han et al., 2014; Gao et al., 2023).

The tectonic evolution of the Precambrian basement in the eastern Hebei terrane during the Neoarchean remains debated. Multiple geodynamic models have been proposed to explain the tectonic setting of Archean metamorphism, including: (1) horizontal tectonic models involving microcontinental block subduction-collision or oceanic slab subduction leading to continent-continent collision (Zhai and Santosh, 2011; Nutman et al., 2011; Fu et al., 2017; Lu et al., 2017; Liu et al., 2018; Kusky et al., 2022; Ning et al., 2023; Wu et al., 2022); (2) Vertical tectonic models associated with mantle plume activity (Zhao et al., 1998; Geng et al., 2006; Yang et al., 2008; Zhao and Zhai, 2013; Kwan et al., 2016); and (3) Archean-specific vertical tectonic regimes associated with sagduction (Duan et al., 2017; Liu et al., 2024; Yu et al., 2022; Li et al., 2024). By integrating the metamorphic evolution and geochronological results obtained from the Luanxian supracrustal rocks with previous studies, we propose that a sagduction model dominated the Neoarchean tectonic regime of the eastern Hebei terrane, evidenced by: (1) Despite multiple phases of metamorphic and deformational overprinting, the Archean dome-and-keel architecture preserved in the eastern Hebei terrane, which distinctly differs from the linear structural patterns of Phanerozoic orogenic belts (Zhao et al., 1999; 2012; Liu et al., 2018; 2024; Zhao et al., 2021; Yu et al., 2022; Xu et al., 2022); (2) Plutonic intrusions (e.g., TTG gneisses, K-rich granites, and quartz diorites) are widespread and coeval, emplaced within a narrow age range (Geng et al., 2006; Yang et al., 2008; Nutman et al., 2011; Wan et al., 2015); (3) The depositional age of supracrustal protoliths shows approximate temporal coincidence with regional magmatism and metamorphism (Zhao and Zhai, 2013; Wan et al., 2015; Duan et al., 2017; Liu et al., 2020; 2024; Zhao et al., 2021; Chen et al., 2024); (4) The Shuangshanzi-Lulong supracrustal belt, located within the Qian’an and Anziling gneiss domes, exhibits a synformal structure with subvertical lineations and sinistral ductile shear zones, indicating downward sag relative to TTG gneisses (Liu et al., 2017; Zhao et al., 2021); (5) The Luanxian Group supracrustal rocks underwent amphibolite-facies metamorphism with a clockwise P-T path and a geothermal gradient of 24°C/km, analogous to the evolution of supracrustal sequences in the Pilbara craton (Figure 7) (François et al., 2014; Cutts et al., 2014), which was attributed to sagduction according to numerical simulation results (François et al., 2014; Sizova et al., 2018; Yu et al., 2022; Liu et al., 2024).