The quartz diorite host rocks (A15.2, A17.1ah, A17.3b, A17-7-HA, A17-14-h, O27b, O52B) were collected from the surface and the drill core. The rocks are coarse grained and granular with the surface samples tending to be more altered than those from the drill core. The rocks are composed of plagioclase (45–50 vol%), amphibole (∼35 vol%), quartz (5–15 vol%), biotite (≤5 vol%), and Fe-Ti oxide minerals (≤5 vol%) with accessory amounts (≤1 vol% each) of apatite, zircon, and titanite. Plagioclase is euhedral to subhedral and with straight crystal edges (Figures 4A, B). The larger plagioclase crystals tend to have oscillatory zonation, but the smaller crystals tend to display polysynthetic twinning with a few showing albite twinning. All plagioclase crystals are altered by saussurite, although the extent of alteration can range significantly between different samples. Amphibole (hornblende) is the primary mafic silicate mineral in the rock and is euhedral to anhedral (Figure 4A). Similar to the plagioclase crystals, there are larger (1–3 mm) euhedral crystals (∼5 vol%), but most are smaller (<1 mm) and interstitial to the plagioclase. Most of the amphibole crystals are altered by chlorite and/or biotite. The quartz crystals are subrounded and interstitial to the plagioclase and amphibole. The amount of quartz can vary between samples and is typically 5–15 vol%. Euhedral to subhedral biotite is common but its abundance can vary between samples (≤5 vol%). The biotite crystals are tan brown to brown in colour with some crystals having zircon inclusions. Under reflected light, it appears that nearly all of the opaque minerals are magnetite and ilmenite as sulphide minerals were not identified. The opaque minerals are euhedral to subhedral with most having cubic or subrounded shapes (Figure 4A).

The diorite rocks (host = A17-8-h, O05c, O20, O27a; enclaves = A17.4b, A17-7-en, A17-8-EN, A17-14-h) were collected from the surface exposure and the drill core. The rocks are phanerocrystalline with some samples having medium to coarse grained textures whereas others are medium to fine grained. In general, the surface samples tend to be more altered than those from the drill core. The mineralogy and textures of the diorites are similar to the quartz diorites, but they have less quartz and biotite (Figures 4C, D). The rocks are primarily composed of plagioclase (50–55 vol%), amphibole (35–40 vol%), and opaque (Fe-Ti oxide minerals) minerals (∼5 vol%) with accessory amounts (≤1 vol% each) of quartz, biotite, apatite, zircon, and titanite.

The mafic rocks (A15.1b, A17.1b, A17.9, A17.10b) have similar textures and mineralogy and are relatively fresh, but show signs of zeolite facies alteration (Figures 4E, F). The rocks are porphyritic and seriate and composed mostly of plagioclase (50–55 vol%) and clinopyroxene (35–40 vol%) with subordinate amounts of ilmenite/magnetite (5–10 vol%), and olivine (∼5 vol%). The plagioclase crystals are mostly euhedral and have lath shapes of similar size. The clinopyroxene (augite) crystals are light brown in colour, subhedral to euhedral in shape. The olivine crystals are euhedral to anhedral (sub-round) and mostly appear as larger phenocrysts. The opaque minerals are euhedral to subhedral with cubic shapes. The basaltic rocks are likely dykes rather than xenoliths as the chilled margin with the host rock was observed (Figures 4G, H). The contact between the quartz diorite host and the basalt is sharp, but there appears to be reaction/transition zone as the contact region changes from brown to dark brown (Figure 4G). The transition zone is aphyric with occasional xenocrysts of quartz and feldspar from the host rock. Farther from the contact, the xenocrysts are absent and the basaltic rock is porphyritic with euhedral to subhedral phenocrysts of olivine and plagioclase microlites (hyalopilitic).

In situ U-Pb and Lu-Hf isotopic analyses presented in this study were performed using a Nu Plasma HR multi-collector inductively coupled plasma mass spectrometer (MC-ICPMS; Nu Instruments, UK) equipped with an ArF excimer 193 nm laser ablation system (RESOlution M-50, Resonetics LLC, USA), housed at the Department of Earth Sciences, The University of Hong Kong, HKSAR. A modified collector block of the MC-ICPMS contains 12 Faraday collectors and 4 ion counting detectors dispersed on the low mass side of the array, allowing simultaneous acquisition of ion signals ranging from mass ²⁰⁴Pb to ²³⁸U. A spot diameter of 30 µm, pulse rate of 6 Hz and energy density of 15 J/cm² were used. Ablation time was 40 s, resulting in pits 20–30 microns deep. Instrument parameters can be referred to Xia et al. (2011). The standard zircon 91500 (Wiedenbeck et al., 1995) and GJ-1 (Jackson et al., 2004) were used for calibration. Off-line data reduction was performed by the software ICPMSDataCal. Version 7.2 (Liu et al., 2010). Ages were calculated using ISOPLOT/Excel version 3.6 (Ludwig, 2008). The results are presented in Supplementary Tables S1, S2 where the isotopic ratios and ages are given with 1 sigma error.

The zircon Hf isotopes were analyzed on the same spots as those with a concordant U-Pb age. The Hf isotopic data reported in this study were obtained with a beam diameter of 55 μm, pulse rate of 6 Hz and energy density of 15 J/cm². Each analytical spot was subjected to 40 ablation cycles, resulting in pits 30–40 μm deep. Atomic masses 172 to 179 were simultaneously measured in static-collection mode. Isobaric interference of ¹⁷⁶Yb on ¹⁷⁶Hf was corrected against the ¹⁷⁶Yb/¹⁷²Yb ratio of 0.5886 (Chu et al., 2002). Interference of ¹⁷⁶Lu on ¹⁷⁶Hf was corrected by measuring the intensity of the interference-free ¹⁷⁵Lu isotope and using a recommended ¹⁷⁶Lu/¹⁷⁵Lu ratio of 0.02655 (Machado and Simonetti, 2001). External calibration was made by measuring zircon standard 91,500 (¹⁷⁶Hf/¹⁷⁷Hf = 0.282275 ± 0.000011) for the unknown samples during the analyses to evaluate the reliability of the analytical data and GJ-1 has been used for monitoring the data quality which yielded a weighted mean ¹⁷⁶Hf/¹⁷⁷Hf ratio of 0.282008 ± 0.000001 (Supplementary Table S3).

The measured ¹⁷⁶Lu/¹⁷⁷Hf and ¹⁷⁶Lu/¹⁷⁷Hf ratios and a¹⁷⁶Lu decay constant of 1.865 x 10⁻¹¹a⁻¹ as reported by Scherer et al. (2001) were used to calculate the initial ¹⁷⁶Hf/¹⁷⁷Hf ratios. Calculations of ε_Hf(t) values were based on the chondritic values of ¹⁷⁶Hf/¹⁷⁷Hf and ¹⁷⁶Lu/¹⁷⁷Hf as reported by Blichert-Toft and Albarède (1997). The mantle extraction model age (T_HfDM) was calculated using the measured ¹⁷⁶Lu/¹⁷⁷Hf of the zircon, but this only provides a minimum age for the source material of the magma from which the zircon crystallized. Therefore, “crustal” model ages T_Hfc were calculated (Supplementary Table S3), which assume that the parental magma of the zircons was produced from average continental crust, but was ultimately derived from the depleted mantle. A ratio of 0.015 most realistically reflects the ¹⁷⁶Lu/¹⁷⁷Hf ratio of our samples.

Three grams (3.0000 ± 0.0005 g) of each sample was added to a ceramic crucible of known mass (+lid) and placed in an oven at 105°C for 3 h. After the initial heating step, the powders were cooled to ambient temperature inside a desiccator for ∼15 min before they were weighed. The samples were then baked to peak temperature at 900°C (held for 10 min) in a high temperature furnace before cooling. The samples were then cooled in a desiccator until reaching ambient temperature. Each sample was weighed for a final time and the loss on ignition was calculated from the masses obtained from the low and high temperature cycles. Six grams (6.0000 ± 0.0005 g) of lithium metaborate was added to each sample (0.6000 ± 0.0005 g) at a ratio of 10:1 and fused to produce a glass disc using a Claisse M4 fluxer. The major elements were measured by wave-length dispersive X-ray fluorescence using Panalytical Axios^mAX at the XRF Laboratory, Department of Earth Sciences, National Taiwan Normal University. Standard reference materials measured during the study include: BIR-1a, BCR-2, and AVG-2a (Supplementary Table S4).

Trace elements were measured using an Agilent 7500cx inductively coupled plasma mass spectrometer (ICP-MS) at National Taiwan University, Taipei, Taiwan. Approximately 50–100 mg of each sample powder was dissolved in a Teflon beaker using a combination of HF, HNO₃, and HCl. Each sample was heated in closed beakers with HF and HNO₃ for at least 48 h and then dried. After drying, 2 mL of 6 N HCl was added to each sample and then left to dry. This step was followed by the addition of 2 mL of 1 N HCl to each sample then the solution was centrifuged. The supernatant was extracted and then placed into a new beaker. If solid residue remained in the beaker after extraction, then the procedure was repeated until the powder was fully digested. Samples were diluted using 2% HNO₃ and an Rh and Bi spike was added for the internal standard. The standard reference materials measured for the trace elements of this study are AGV-2 (andesite), BCR-2 (basalt), and DNC-1 (dolerite) (Supplementary Table S4).

Whole rock Sr and Nd isotopes were measured at the Institute of Earth Sciences, Academia Sinica, Taipei (Supplementary Table S5). Approximately 75 mg of sample powder was dissolved using a mixture of HF and HNO₃ at 120°C for 48 h, then taken to dryness followed by dissolution in 2 mL 4 N HCl. Separation and purification of Sr and Nd were achieved using a 3-column technique. The first column was contained 0.2 mL Sr spec resin (manufactured by Eichron Industries, Inc.) to collect the Sr fraction using HNO₃ as eluent. The elution was then passed through the second column containing 2.5 mL cation exchange resin (Bio Rad AG50W-X8, 100–200 mesh) to isolate the rare Earth elements using HCl as eluent. The third column used for Nd separation contained 1 mL Ln-B25-A (Eichrom) resin, covered on top by a thin layer of anion exchange resin (Bio Rad AG1-X8, 200–400 mesh), using HCl as eluent.

Strontium was loaded on a single Ta filament and compositional analysis was performed using 7-collector Finnigan MAT-262 thermal ionization mass spectrometer in dynamic mode. Neodymium loaded on a double Re filament and the isotopic measurement was analyzed using 9-collector Thermo Fisher Scientific Triton thermal ionization mass spectrometer with static mode for automatic run. The isotopic ratios were corrected for mass fractionation by normalizing to ⁸⁶Sr/⁸⁸Sr = 0.1194 and ¹⁴⁶Nd/¹⁴⁴Nd = 0.7219. At the time of analysis, the NBS-987 Sr standard yielded an average value of ⁸⁷Sr/⁸⁶Sr = 0.710249 with a long-term reproducibility of 0.000045 (2σ) and the JMC Nd standard gave an average of ¹⁴³Nd/¹⁴⁴Nd = 0.511813 with a long-term reproducibility of 0.000009.

Twenty-six zircons were analyzed from sample O05c (diorite). Zircon grains range in size from ∼50 μm to ∼300 µm in length and have euhedral (prismatic to equant) to subhedral (fragmented to sub-round) shapes. Nearly all zircons show igneous (oscillatory) zonation whereas very few have core-rim structures (Corfu et al., 2003). All zircons yielded ages that were 92% concordant or better, however we only select zircons with ≥97% concordance to interpret their emplacement age (Supplementary Table S1). The ²⁰⁶Pb/²³⁸U ages revealed a bimodal distribution of ∼111 Ma and ∼107 Ma. Five zircons yielded an intercept age of 111.3 ± 1.8 Ma and weighted-mean ²⁰⁶Pb/²³⁸U age of 111.3 ± 0.5 Ma which we interpret as an inheritance age (Figure 5A). The majority of zircons (12) yielded an intercept age of 107.4 ± 1.0 Ma and weighted-mean ²⁰⁶Pb/²³⁸U age of 107.5 ± 0.3 Ma which we interpret as the emplacement age (Figure 5B). The older group of zircons (∼111 Ma) do not show significantly different internal structures (i.e., core-rim structures) nor are their crystal shapes significantly different from the younger (∼107 Ma) group of zircons.

Thirty-eight zircons were analyzed from sample O20 (diorite), however only twenty-six zircons with ≥97% concordance are considered for the emplacement ages (Supplementary Table S2). Zircon grains range in size from ∼40 μm to ∼400 µm in length and have euhedral (prismatic to equant) to subhedral (fragmented to sub-round) shapes. The zircons are morphological similar to those from sample O05c. The ²⁰⁶Pb/²³⁸U ages range from ∼106 Ma to ∼577 Ma with the majority (9) of ages around ∼112 Ma. The nine zircons yielded an intercept age of 112.6 ± 1.4 Ma with a weighted-mean age of 112.4 ± 0.4 (Figure 5C). Four zircons with ≥97% concordance yielded an intercept age of 109.2 ± 1.3 Ma and weighted-mean ²⁰⁶Pb/²³⁸U age of 109.1 ± 0.4 Ma which we interpret as the emplacement age as they are within the uncertainty of the ages of O05c (Figure 5D). The inherited zircons with ≥97% concordance are mostly Carboniferous to Permian (258–335 Ma), but there are single zircon ages of ∼106 Ma, ∼114 Ma, ∼118 Ma, and ∼577 Ma. Much like sample O20, there does not appear to be significantly different internal structures or crystal shapes between the two groups of Cretaceous zircons. However, the older inherited zircons (i.e., >250 Ma) tend to have irregular to rounded shapes.

Hafnium isotopes were measured on the same zircon spot locations as the U-Pb ages of samples O20 and O05c. Not all zircons could be analyzed for Hf isotopes due to the limited amount of material left after ablation for U-Pb dating. The dataset can be found in the Supplementary Table S3. The depleted mantle lines in Figure 6 are defined by present-day ¹⁷⁶Hf/¹⁷⁷Hf = 0.28325 and ¹⁷⁶Lu/¹⁷⁷Hf = 0.0384 (Griffin et al., 2000).

The magmatic zircons (12) from sample O20 yielded initial ¹⁷⁶Hf/¹⁷⁷Hf(t) ratios from 0.283048 to 0.283186 with ε_Hf(t) values ranging from +12.1 to +17.0. The meaningful T_DM1 ages range from 123 Ma to 291 Ma whereas the T_DM2 ages range from 141 Ma to 543 Ma. In comparison, the inherited (296–335 Ma) zircons (5) have initial ¹⁷⁶Hf/¹⁷⁷Hf(t) ratios from 0.282675 to 0.283025 and ε_Hf(t) values ranging from +3.9 to +15.5. The T_DM1 ages range from 313 Ma to 813 Ma whereas the T_DM2 ages range from 341 Ma to 1,490 Ma.

Twenty-two magmatic zircons from sample O05c yielded initial ¹⁷⁶Hf/¹⁷⁷Hf(t) ratios from 0.282904 to 0.283113 with corresponding ε_Hf(t) values from +7.0 to +14.4. The T_DM1 ages range from 194 Ma to 493 Ma whereas the T_DM2 ages range from 319 Ma to 1,040 Ma. There is one inherited zircon (240 Ma). It has an initial ¹⁷⁶Hf/¹⁷⁷Hf(t) ratio of 0.282657 and ε_Hf(t) value of +1.2 and T_DM1 and T_DM2 ages of 830 Ma and 1,688 Ma.

The SiO₂ contents of the host rocks range from ∼55.7 to ∼71.6 wt% with nearly all ANCK values (mol. Al₂O3/[mol. CaO+Na₂O+K₂O]) between 0.87 and 1.01 and alkali-alumina (mol. [Na+K]/mol. Al) values <1 (Figures 7A, B). According to the scheme of Frost et al. (2001), the rocks classify as magnesian (FeOt/FeOt+MgO = 0.65–0.71) and calcic (Figures 7C, D). Magnesian, metaluminous, calcic rocks are considered to be typical of the outboard portions of Cordilleran Batholiths (i.e., volcanic-arc granites). Sample A15.2 has very low Na₂O and K₂O contents and consequently has a very high ANCK value (∼1.9). However, this sample is altered and it is likely that the alkali metals were mobilized (L.O.I. = 3.21 wt%). Nonetheless, for the exception of A15.2, all rocks are sodic (K₂O/Na₂O = 0.28–0.71) and their Mg# ([mol Mg/(mol Mg+tFe²⁺)]× 100) are 42.1–49.4 (Figures 7E, F). The loss on ignition (L.O.I.) values, for the exception of A15.2, are ≤0.51 wt% and indicate that the rocks did not undergo significant post-emplacement hydrothermal/deuteric alteration.

The host rocks have uniformly low concentrations of transition metals (Sc = 5.4–16.8 ppm), V = 93.6–170.6 ppm, Cr = 6–19 ppm, Co = 8.5–16.3 ppm, Ni = 4–8 ppm, Cu = 4.9–34.7 ppm, Zn = 16.7–55.3 ppm). The large ion lithophile element (LILE) concentrations are more variable (Rb = 0.9–84.9 ppm, Sr = 77–679 ppm, Cs = 0.12–2.3 ppm, Ba = 34–538 ppm). Similar to the LILE, the high field strength elements are variable (Zr = 22–108 ppm, Nb = 3.1–4.8 ppm, Y = 9.3–23.5 ppm, Hf = 0.78–2.68 ppm, Ta = 0.27–0.37 ppm, Th = 2.6–6.4 ppm, U = 1.1–1.9 ppm). The mid-ocean ridge basalt normalized incompatible patterns of the rocks are broadly similar, although there is some variability in the Rb−Ba and Hf−Zr concentrations (Figure 8A). All rocks have similar chondrite normalized rare Earth element patterns (Figure 8B). The rocks are light rare Earth element (REE) enriched with high La_N/Sm_N (1.9–3.3) and flat to bowl shaped middle to heavy REEs (Gd_N/Yb_N = 1.1–1.6) patterns. The Eu/Eu* values [2*Eu_N/(Sm_N+Gd_N)] are 0.93–1.09.

The four dioritic enclaves have intermediate compositions with SiO₂ ranging from 54.8–57.6 wt% and Na₂O+K₂O wt% equal to 5.2–5.6 wt% (Figure 7A). The enclaves are metaluminous, magnesian, and alkali-calcic (Figures 7B–D). The Fe₂O₃t ranges from 7.5–8.9 wt% and MgO from 3.9–5.0 wt% with the Mg# ranging from 46.6–53.6. The TiO₂ (0.65–0.68 wt%) and MnO (<0.25 wt%) contents are low and the CaO is ∼7.5–7.8 wt%. The transition metals decrease with increasing SiO₂ contents, although some elements have limited variability (Sc = 14.1 and 21.9 ppm, V = 190 and 195 ppm, Cr = 18 and 38 ppm, Co = 16 and 17.5 ppm, Ni = 10 and 12 ppm, Cu = 22.4, and 38.4 ppm, Zn = 38.9–44.5 ppm). The mid-ocean ridge ocean basalt normalized incompatible element patterns are similar to the host rocks with enriched patterns of Cs, Rb, Ba, Th, and U with a depletion of Nb and relatively flat La−Nd patterns and flat but lower concentrations of Hf−Y with a positive anomaly of Sm (Figure 8C). The enclaves have light REE-enriched patterns with (La/Yb)_N and (La/Sm)_N ratios of 4.9–5.0 and 2.4. The rocks do not have negative Eu anomalies with Eu/Eu* ratios from 0.97–0.98 (Figure 8D). Furthermore, the rocks tend to have bowl shape patterns as the middle rare Earth elements (Dy−Er) are flat and the heavy rare Earth elements (Tm−Lu) are slightly higher.

The four basaltic rocks have SiO₂ ranging from 48.7–50.6 wt% and Na₂O+K₂O wt% equal to 4.2–5.4 wt% (Figure 7A). There is a distinct SiO₂ gap between the basaltic (∼50.5 wt%) rocks and dioritic (∼54.8 wt%) enclaves (Figures 7A, F). The basalts are mildly alkaline with alkalis (Na₂O+K₂O) increasing with increasing SiO₂ contents. The Fe₂O₃t ranges from 9.6–11.2 wt% and MgO from 4.9–9.1 wt% with the Mg# ranging from 50.1–61.7. The TiO₂ (1.5–1.6 wt%) and MnO (<0.20 t.%) contents are low. The Al₂O₃ ranges from 15.3–18.2 wt%, and CaO from 9.4–10.6 wt%. The transition metals generally decrease with increasing SiO₂ contents, although some elements have limited variability (Sc = 7.9–10.1 ppm, V = 219–224 ppm, Cr = 16–354 ppm, Co = 28.7–48.9 ppm, Ni = 24–205 ppm, Cu = 45.1–58.6 ppm, Zn = 81.5–111.7 ppm). The mid-ocean ridge ocean basalt normalized incompatible element patterns show low Cs and Rb concentrations and declining concentrations from Ba to Y (Figure 8E). The basalts have chondrite normalized light REE-enriched patterns with (La/Yb)_N and (La/Sm)_N ratios between 11.2 and 13.2 and 2.7–3.1. The rocks do not have significant negative Eu anomalies with Eu/Eu* ratios from 0.93–0.96 (Figure 8F).

Five samples were selected for Sr and Nd isotopic analysis (Supplementary Table S5). The initial isotopic ratios of the diorites and quartz diorites were calculated based on the U-Pb ages of this study (107 Ma) whereas the initial isotopic ratios of the basaltic rocks are thought to be contemporaneous with the adjacent Neogene−Quaternary volcanic rocks and an age of 10 Ma was assumed (Nguyen and Flower, 1998). The initial ⁸⁷Sr/⁸⁶Sr ratios of the granitic rocks range from 0.704313 to 0.707681 (Figure 9). The initial ¹⁴³Nd/¹⁴⁴Nd ratios range from 0.512660 to 0.512750 for the granitic samples. Their ε_Nd(t) values, using ¹⁴⁷Sm/¹⁴⁴Nd of 0.1967 and a CHUR_today value of 0.512638, correspond to +3.1– +4.9 (Figure 9). The depleted mantle model ages (T_DM) range from 598 Ma to 847 Ma. The lone basalt sample has an initial ⁸⁷Sr/⁸⁶Sr ratio of 0.703978, an ε_Nd(t) value of +4.2, and T_DM age of 513 Ma.

The zircon U-Pb geochronology results of the surface rocks yielded Albian ages (O05c = 107.5 ± 0.3 Ma; O20 = 109.1 ± 0.4 Ma). Moreover, sample O20 has inherited zircons that range in age from Late Ediacaran (∼577 Ma) to Early Cretaceous (∼118 Ma) whereas sample O05c has one early Triassic zircon (∼240 Ma). The zircon ages of the diorites from this study are contemporaneous with Early Cretaceous granitic magmatism throughout the Dalat-Kratie Fold Belt of Vietnam and Cambodia (Nguyen et al., 2004a; Shellnutt et al., 2013; Breitfeld et al., 2020; Hennig-Breitfeld et al., 2021; Kasahara et al., 2021; Nong et al., 2021; Nong et al., 2022). Specifically, they correlate to a period of subduction-related (130–90 Ma) magmatism along the western margin of the Paleo-Pacific Ocean (Taylor and Hayes, 1983; Metcalfe, 2006; Yang, 2013; Hennig-Breitfeld et al., 2021; Waight et al., 2021).

Farther north along the SE coast of China, the Early Cretaceous period of arc-magmatism is referred to as the Late Yanshanian Orogen and is typified by I-type granite magmatism with subordinate S-type granite magmatism (Zhou et al., 2006; Li et al., 2012; Li et al., 2014; Dong et al., 2018; Shellnutt et al., 2020a; Suga and Yeh, 2020). From a geodynamic point of view, the Early Cretaceous Eurasian margin arc-magmatism is related to subduction of Paleo-Pacific oceanic crust (Maruyama et al., 1997; Yang, 2013; Wu et al., 2022).

Inherited zircons are more common in sample O20 whereas there was only one identified in O05c (∼240 Ma). Sample O20 has the most inherited zircons (22) and a clear Paleozoic population (335–296 Ma) and one Late Ediacaran zircon (∼577 Ma). There are eleven inherited zircons with a slightly older age range (112–118 Ma) than the weighted-mean ages of the diorite. We interpret these inherited zircons as being derived from an older period of spatially associated subduction-related magmatism as they have similar Hf isotopic values (ε_Hf(t) = +13.2– +16.8) as the emplacement age zircons (ε_Hf(t) = +12.1– +17.0). Moreover, the inherited Early Cretaceous (112–118 Ma) ages overlap with reported rock ages throughout the Dalat-Kratie Fold Belt of Vietnam and Cambodia (Nguyen et al., 2004a; Shellnutt et al., 2013; Hennig-Breitfeld et al., 2021; Kasahara et al., 2021; Nong et al., 2021). The Late Paleozoic (335–296 Ma) inherited zircons are contemporaneous with Middle Carboniferous (336 ± 3 Ma) to Permian (267 ± 3 Ma) mafic and silicic rocks known throughout Indochina. The magmatic rocks may be related to an earlier period of subduction of the Paleotethys and back-arc extension prior to the amalgamation between the Sukhothai and Indochina terranes or subduction of the Paleo-Pacific plate (Hall and Sevastjanova, 2012; Burrett et al., 2014; 2021; Wang et al., 2018; Cheng et al., 2019; Kasahara et al., 2021; Shi et al., 2021; Waight et al., 2021; Breitfeld et al., 2022).

The Triassic zircon of O05c is contemporaneous with rocks in Cambodia as Cheng et al. (2019), Kasahara et al. (2021) and Waight et al. (2021) reported zircon U-Pb ages from 201.88 ± 0.36 Ma to 238.21 ± 0.31 Ma. The Triassic magmatism in Cambodia is interpreted to be related to subduction and back-arc extension associated with Paleotethys subduction along the western margin of the amalgamated Sukhothai-Indochina terranes (Sone and Metcalfe, 2008; Kasahara et al., 2021; Waight et al., 2021). Moreover, there is Triassic magmatism throughout Vietnam that may be related to subduction of the Paleo-Pacific plate (Hennig-Breitfeld et al., 2021).

Many continental and island arcs show evidence of magmatic migration over time (e.g., Lesser Antilles, Andes, Sunda) that is thought to be related to changes in the subducting plate dynamics (Kay et al., 2005; He and Xu, 2012; Allen et al., 2019; Lai et al., 2021). The Late Mesozoic granites of eastern China are an example of secular variation of granitic batholiths in a convergent margin regime as magmatism changed from one of sinistral strike-slip compression during the Jurassic to an active continental margin during the Early Cretaceous (Zhou et al., 2006; Dong et al., 2018). The Cretaceous granitic batholiths of eastern China show chemo-temporal transition from subduction-related Cordilleran-type batholiths during the Early Cretaceous (130–100 Ma) to extension-related A-type granitic batholiths during the Late Cretaceous (<90 Ma) (Zhou and Li, 2000; Zhou et al., 2006; Li et al., 2012). Correlative with the Late Yanshanian granites of eastern China are the Early Cretaceous granitic batholiths of the southern Indochina terrane (Nguyen et al., 2004a; Nguyen et al., 2004b; Shellnutt et al., 2013).

The reported ages of the Dalat-Kratie granites are mostly between ∼120 Ma and ∼80 Ma and show a chemical transition from Cordilleran I-type granites to post-collisional (A-type) granites over time (Nguyen et al., 2004a; Nguyen et al., 2004b; Shellnutt et al., 2013; Hennig-Breitfeld et al., 2021). Although there is a compositional transition from I-type granite to A-type granite, there does not appear to be a spatial correlation as the older I-type granites (120-115) geographically overlap with the younger I-type granites (110-90 Ma) and the A-type granites (<90 Ma) (Nguyen et al., 2004a; Nguyen et al., 2004b; Shellnutt et al., 2013; Cheng et al., 2019; Kasahara et al., 2021; Hennig-Breitfeld et al., 2021; Waight et al., 2021; Nong et al., 2022). However, there is a distinct change in the Sr-Nd isotopic compositions of the I-type granites in Cambodia compared to the granites of similar age in Vietnam (Figure 9).

The precise reason for the isotopic change is not clear, but it must be related to the parental magma source. Consequently, there are two possibilities: 1) there is a change in the lithospheric composition of each region, or 2) there were differences in the amount of isotopically enriched components in the genesis of the parental magmas. If the parental magmas of the I-type granites of each region are derived primarily by lithospheric melting then the western region must be isotopically juvenile compared to the eastern region. This could imply that there is a heretofore unseen lithotectonic boundary between the two regions (c.f., Dijkstra and Hatch, 2018). Alternatively, it could reflect differences in the amount of enriched material (e.g., fluids, melts) originating either from subducted sediment or crustal melts that is incorporated into the parental magmas whether they are derived from the mantle or subducted slab. That is, there was more enriched material involved in the genesis of the granitoids from Vietnam than the granitoids from Cambodia.

We cannot be certain which of the two scenarios is more likely, although we tend to favour the possibility that there is a domain boundary between eastern Cambodia and southern Vietnam (Figure 14). The western region of the Dalat-Kratie belt in Vietnam and Cambodia is referred to as the Srepok-Tay Nam Bo orogenic belt (c.f., Hennig-Brietfeld et al., 2021). Consequently, it is possible that the Srepok-Tay Nam Bo orogenic belt and the Dalat-Kratie belt are compositionally and tectonically distinct regions of the Indochina terrane. Our preferred interpretation is based on the fact that the Carboniferous-Permian (343–253 Ma) inherited zircons from the Late Cretaceous (86.6 ± 1.6 Ma) Ankroet “A-type” granite of Vietnam have negative ε_Hf(t) values (−10.2– −3.3; Shellnutt et al., 2013) whereas the Carboniferous-Permian (335–296 Ma) inherited zircons from the Snoul pluton have positive ε_Hf(t) values (+3.9− +15.5). Furthermore, the single Triassic (249 ± 5 Ma, 1σ) inherited zircon from Ankroet has a highly negative ε_Hf(t) value (−17.2) whereas in the Snoul pluton the Early Triassic (240 Ma) inherited zircon has near chondritic value ε_Hf(t) values (+1.2). Thus, the Hf isotopes of the inherited Paleozoic zircons and the whole rock Sr-Nd isotopes indicate that the isotopic differences between the two regions are likely to be an intrinsic feature of the Indochina lithosphere that predates the Cretaceous (e.g., Shellnutt et al., 2020b). Therefore, it possible that the lithosphere of the southern Indochina terrane is composed of distinct lithotectonic domains.