Compositional variations of Neoproterozoic I-type granitoids in response to slab breakoff in the western Yangtze block, South China

Abstract

The Neoproterozoic western Yangtze block is a vital area for understanding geodynamics of the Rodinia supercontinent. In order to explore temporal and compositional variations of I-type granitoids in this region, we present zircon U-Pb-Hf isotopes and bulk rock geochemical data of the Shimian monzogranite and Kangding tonalite. They yield emplacement ages of 827 ± 5 Ma and 757 ± 3 Ma with ε_Hf(t) values ranging from 6.8 to –1.1 and 9.6 to 3.6, respectively. Typical peraluminous and high-temperature minerals, such as garnet and pyroxene, are scarce. The characteristics of low A/CNK, Ga/Al and FeO^T/MgO ratios, as well as low zircon saturation temperatures (770–660°C) can be identified from the granitoids, suggest a typical I-type granite origin. They have variable Y/Nb ratios of 6.18–1.76, Mg^# values of 44.3–14.2, and ε_Hf(t) values of 9.6 to –1.1, which can be explained by a heterogeneous source primarily comprising juvenile crust with minor ancient crustal components. Elemental correlations are indicative of fractional crystallization, such as Fe-Ti oxides and plagioclase. In comparison, regional I-type granitoids have geochemical diversities which are bounded by ca. 780 Ma, e.g., contrasting SiO₂ contents (average 71.78 wt.% vs. 63.79 wt.%), Y/Nb ratios (average 3.91 vs. 3.49), Rb/Sr ratios (average 1.52 vs. 0.13), Mg^# values (average 27.9 vs. 43.6) and ε_Hf(t) values (average 4.84 vs. 7.76). All these evidences suggest a different participation of juvenile crust in their source region before and after ca. 780 Ma. Combined with previous studies, we propose an arc affinity for the Neoproterozoic granitoids in this region, and a slab breakoff on ca. 780 Ma responsible for their compositional variations.

Introduction

I-type granitoids, defined as granitic rocks derived from original igneous rocks, are one of the most universal continental granite types in terms of their forming mechanism (Chappell, 1999). They commonly display wide variations of chemical compositions, e.g., SiO₂ ranging from <65 wt.% to >75 wt.% (Clemens et al., 2011). Multiple factors can lead to these chemical characteristics, including distinct source rocks (Lai et al., 2015), partial melting condition (Zhao et al., 2015), anatectic temperature (Gao et al., 2016) and magma mixing (Yang et al., 2007). Previously proposed forming mechanisms can be classified into three categories: (1) fractional crystallization of mantle-derived mafic melts (Wyborn et al., 1987; Clemens et al., 2016), (2) partial melting of pre-existing crustal igneous rocks (Barbarin, 1996; Gao et al., 2016), and (3) mixing of mantle-crust compositions (Castro et al., 1991; Weidendorfer et al., 2014). Hence, deciphering the geochemical diversities and petrogenesis of I-type granitoids can yield critical information on their forming conditions and broader implications, such as crust-mantle interactions and geodynamic processes (Kemp et al., 2007).

The South China craton comprises the northwestern Yangtze block and southeastern Cathaysia block (Fig. 1a), separated by the Neoproterozoic Jiangnan orogen (Li et al., 2008). Its assembly and breakup are contentious in modern geosciences (e.g., Li et al., 2002, 2008; Cawood et al., 2013; Zhu et al., 2021). Neoproterozoic igneous rocks are widely distributed in the western Yangtze block (WYB) and its periphery, with ages ranging from 860750 Ma (e.g., Zhou et al., 2002; Zhao et al., 2008a, 2018; Du et al., 2014; Chen et al., 2015). Among these igneous suites, calc-alkaline I-type granitoids are widely spreaded (Zhao et al., 2018); however, few studies have tested the geochemical diversities and their link with temporal variation.

Fig. 1.

Location (a) and geological map (b) of the WYB, modified after Du et al. (2014).

In this paper, we examined the Neoproterozoic I-type granitoids form the South China block, aiming to unravel their petrogenesis and geodynamic implications. Zircon U-Pb-Hf isotopes and bulk rock chemical data are presented for ca. 827 Ma Shimian monzogranite and ca. 757 Ma Kangding tonalite (Fig. 1b). The data was further compared with typical Neoproterozoic I-type granitoids in the WYB. Integrating results from previous studies, this study further explores the tectonic evolution and geodynamic implications.

Geological Setting and Samples

Regional geology

The South China craton is tectonically divided into the Yangtze and Cathaysia blocks, which are separated by ca. 830 Ma Jiangnan orogenic belt (Zhao et al., 2011). The Yangtze block is situated in the northwestern South China craton, and is segregated from the North China craton by the Qinling-Dabie orogenic belt in the north, and the Tibetan Plateau by the Longmenshan fault in the west (Fig. 1a). This block is underlain by Mesoproterozoic arenaceous to argillaceous metasedimentary packages. The Archean to Paleoproterozoic crystalline basements are sporadically distributed. Particularly, the oldest Kongling complex formed at ca. 3.45 Ga contains TTG gneisses, metasedimentary rocks and amphibolites (Guo et al., 2014). The Neoproterozoic strata include volcanic-sedimentary sequences, with 860–740 Ma igneous rocks mainly distributed along the Yangtze block margins (e.g., Zhou et al., 2002, 2006).

The WYB consists of Proterozoic sequences including metamorphic complexes, metasedimentary strata and volcanic successions. The metamorphic complexes, such as the Kangding complex, comprise granitic gneisses and were formerly considered as Archean to Paleoproterozoic basement (He et al., 1988; BGMR of Sichuan Province, 1991); however, recent geochronological data yield Neoproterozoic ages (860–750 Ma, e.g., Zhou et al., 2002; Zhao et al., 2008b, 2018). The 803 Ma Suxiong volcanic rocks are located in the northern WYB (Li et al., 2002). These volcanic rocks unconformably overly Mesoproterozoic slates, and are conformably covered by Ediacaran to Permian strata (BGMR of Sichuan Province, 1991).

Large amounts of Neoproterozoic igneous rocks are widely distributed in the WYB (Fig. 1b), including granites, granodiorites, tonalites and mafic-ultramafic bodies (Zhao et al., 2018). Among these igneous suits, I-type granitoids are spatially preserved within basement complexes, and were emplaced for about 100 Ma (e.g., Zhou et al., 2002, 2006). This is exemplified by ca. 828 Ma Qinganlin granite (Chen et al., 2015), ca. 790 Ma Shimian monzogranite (Zhao et al., 2008b), ca. 781.1 Ma Dalu granodiorite (Zhu et al., 2019), ca. 779.8 Ma Dalu granite (Zhu et al., 2019), ca. 754 Ma Luding monzodiorite (Lai et al., 2015) and ca. 748 Ma Luding granodiorite (Lai et al., 2015).

Petrography

For detecting compositional variations of Neoproterozoic I-type granitoids with different ages in the WYB, two granitoids near Shimian and Luding were collected in the present study. Their sampling positions are marked in Fig. 1b.

The Shimian samples are medium- to coarse-grained monzogranite. They primarily consist of K-feldspar (~40%, 0.5–3.0 mm), quartz (~30%, 0.3–1.5 mm) and plagioclase (~20%, 0.5–2.0 mm), with minor biotite (~5%, 0.5–1.0 mm) and hornblende (~5%, 0.2–0.8 mm). Commonly, plagioclase grains show polysynthetic twinning while K-feldspar crystals display crossed twinning. They locally enclose biotite and hornblende (Figs. 2a and b).

Fig. 2.

Photographs of the Shimian monzogranite (a–b) and the Kangding tonalite (c–d). Abbreviation: Qtz = quartz, Kfs = K-feldspar, Pl = plagioclase, Bi = biotite, Hbl = hornblende. The size of the coins is about 2 cm.

The Kangding samples are fine- to medium-grained tonalite, and are characterized by gneissic structure. They are dominated by plagioclase with polysynthetic twinning (~45%, 0.5–1.0 mm), quartz (~30%, 0.2–1.0 mm), biotite (~15%, 0.3–0.8 mm), and minor hornblende (~5%, 0.1–0.3 mm) and K-feldspar (~5%, 0.5–1.0 mm). Biotite and quartz grains are commonly subhedral, and quartz crystals show wavy extinction in some cases (Figs. 2c and d).

Analytical Methods

In this study, samples SM-1 and KD-1 (Figs. 2a and c) were chosen for zircon U-Pb-Hf isotopes. All rock samples were powdered to 200 meshes for bulk rock chemical analyses.

Zircon grains were separated by magnetic and density methods, and then mounted in epoxy. Their internal morphology was detected using cathodoluminescence (CL) images. Zircon U-Pb dating was performed at Nanjing FocuMS Technology Co. Ltd, using Teledyne Cetac Technologies Analyte Excite laser-ablation system and Agilent Technologies 7700x quadrupole ICP-MS. Detailed procedures and parameters are given by Dai et al. (2017). Standards 91500 and GJ-1 were analysed every eight spots. They have mean ²⁰⁷Pb/²⁰⁶Pb age of 1079 ± 42 Ma (n = 14, MSWD = 0.15, 95% conf.) and ²⁰⁶Pb/²³⁸U age of 603.6 ± 3.3 Ma (n = 7, MSWD = 0.88, 95% conf.), respectively, within errors given in Wiedenbeck et al. (1995) and Jackson et al. (2004). Raw data were processed by the method of Liu et al. (2010).

Zircon Hf isotopes were also measured at Nanjing FocuMS Technology Co. Ltd, using Nu Instruments Nu Plasma II MC-ICP-MS. They were conducted on the same grains for U-Pb dating. Standards Plešovice and Mud Tank were analysed every eight spots. They yielded average ¹⁷⁶Hf/¹⁷⁷Hf ratios of 0.282478 ± 0.000008 and 0.282515 ± 0.000011, respectively, compatible to recommended ratios in Sláma et al. (2008) and Woodhead and Hergt (2005).

Whole-rock geochemical analyses were carried out in the Analytical Laboratory of ALS Chemex (Guangzhou) Company Limited. Detailed procedures are provided by Dai et al. (2017). Major oxides were analysed by a PANalytical Axios-advance (Axios PW4400) X-ray fluorescence spectrometer, with analytical precision less than 5%. Trace elements were measured by a Perkin-Elmer Elan 9000 ICP-MS, with a precision better than 10%.

Results

Zircon U-Pb geochronology

Dating results of samples SM-1 and KD-1 are presented in Table 1. CL images show euhedral morphology and apparent oscillatory zoning for most zircon grains (Figs. 3a and b). They have lengths up to 130 μm, with length/width ratios of 2:1 and Th/U ratios up to 1.3. This is suggestive of magmatic zircons (Hoskin and Schaltegger, 2003).

Table 1. Zircon U-Pb-Hf isotope data of the Neoproterotoic Shimian monzogranite and Kangding tonalite

Spots	Th	U	Th/U	207Pb/206Pb	±σ	207Pb/235U	±σ	206Pb/238U	±σ	207Pb/235U Age (Ma)	±σ	206Pb/238U Age (Ma)	±σ	Concordance	176Yb/177Hf	176Lu/177Hf	176Hf/177Hf	±2σ	εHf(t)	TDM1 (Ma)	TDM2 (Ma)
The Shimian monzogranite
SM-1.1	570	1131	0.5	0.0673	0.0010	1.2922	0.0236	0.1395	0.0018	842	10	842	10	99%	0.046278	0.001613	0.282433	0.000010	5.3	1181	1383
SM-1.2	394	686	0.6	0.0666	0.0008	1.2521	0.0159	0.1357	0.0007	824	7	820	4	99%	0.066413	0.001934	0.282293	0.000010	0.2	1392	1707
SM-1.3	418	576	0.7	0.0656	0.0009	1.2448	0.0158	0.1371	0.0007	821	7	828	4	99%	0.079507	0.002237	0.282346	0.000009	1.9	1327	1600
SM-1.4	313	520	0.6	0.0665	0.0008	1.2508	0.0158	0.1360	0.0007	824	7	822	4	99%	0.075738	0.002390	0.282264	0.000010	–1.1	1451	1788
SM-1.5	584	2868	0.2	0.0680	0.0007	1.2826	0.0132	0.1361	0.0008	838	6	823	4	98%	0.060485	0.002363	0.282294	0.000009	0.0	1407	1721
SM-1.6	293	598	0.5	0.0673	0.0009	1.2999	0.0162	0.1398	0.0008	846	7	844	4	99%	0.074996	0.001810	0.282262	0.000011	–0.9	1431	1772
SM-1.7	771	893	0.9	0.0654	0.0008	1.2484	0.0146	0.1378	0.0008	823	7	832	4	98%	0.061811	0.001909	0.282288	0.000009	0.0	1399	1718
SM-1.8	873	718	1.2	0.0669	0.0008	1.2796	0.0157	0.1381	0.0008	837	7	834	4	99%	0.050583	0.001533	0.282475	0.000007	6.8	1118	1286
SM-1.9	402	406	1.0	0.0698	0.0010	1.3308	0.0187	0.1380	0.0008	859	8	833	4	96%	0.050686	0.001666	0.282430	0.000011	5.2	1186	1391
SM-1.10	525	1222	0.4	0.0675	0.0009	1.3082	0.0162	0.1402	0.0007	849	7	846	4	99%	0.023950	0.000852	0.282437	0.000013	5.9	1151	1348
SM-1.11	162	737	0.2	0.0665	0.0008	1.2529	0.0140	0.1364	0.0007	825	6	824	4	99%	0.045071	0.001369	0.282416	0.000008	4.8	1197	1413
SM-1.12	362	522	0.7	0.0667	0.0016	1.2662	0.0296	0.1378	0.0012	831	13	832	7	99%	0.064961	0.001560	0.282336	0.000011	1.9	1317	1598
SM-1.13	220	595	0.37	0.0664	0.0014	1.2317	0.0254	0.1340	0.0012	815	12	811	7	99%	0.020598	0.000748	0.282454	0.000011	6.5	1125	1306
SM-1.14	254	473	0.54	0.0673	0.0012	1.2623	0.0225	0.1355	0.0010	829	10	819	5	98%	0.041905	0.001438	0.282424	0.000012	5.1	1188	1398
SM-1.15	73.9	119	0.62	0.0654	0.0019	1.2223	0.0351	0.1352	0.0011	811	16	817	6	99%	0.029524	0.001064	0.282448	0.000013	6.2	1142	1330
SM-1.16	451	848	0.53	0.0632	0.0010	1.1861	0.0185	0.1353	0.0009	794	9	818	5	97%	0.041764	0.001405	0.282399	0.000013	4.2	1222	1451
SM-1.17	329	601	0.55	0.0633	0.0012	1.1825	0.0222	0.1348	0.0010	792	10	815	6	97%	0.045927	0.001512	0.282409	0.000012	4.5	1211	1433
SM-1.18	236	537	0.44	0.0639	0.0012	1.1991	0.0220	0.1356	0.0010	800	10	820	6	97%	0.039752	0.001357	0.282421	0.000012	5.0	1189	1401
SM-1.19	167	395	0.42	0.0663	0.0012	1.2479	0.0227	0.1359	0.0009	822	10	821	5	99%	0.044596	0.001737	0.282419	0.000013	4.7	1205	1419
SM-1.20	376	1078	0.35	0.0650	0.0009	1.2116	0.0173	0.1347	0.0008	806	8	815	4	98%	0.024060	0.000926	0.282420	0.000011	5.2	1178	1389
The Kangding tonalite
KD-1.1	148	163	0.9	0.0646	0.0015	1.0930	0.0238	0.1236	0.0009	750	12	751	5	99%	0.025300	0.000982	0.282521	0.000016	7.2	1038	1207
KD-1.2	102	121	0.8	0.0652	0.0016	1.1207	0.0260	0.1257	0.0010	763	12	763	6	99%	0.021833	0.000614	0.282546	0.000011	8.3	993	1138
KD-1.3	195	208	0.9	0.0648	0.0012	1.1268	0.0206	0.1261	0.0009	766	10	765	5	99%	0.027500	0.000943	0.282571	0.000012	9.0	966	1092
KD-1.4	172	144	1.2	0.0654	0.0011	1.1223	0.0176	0.1245	0.0007	764	8	756	4	98%	0.037040	0.001180	0.282490	0.000013	6.0	1087	1283
KD-1.5	122	139	0.9	0.0641	0.0011	1.1010	0.0175	0.1248	0.0010	754	8	758	6	99%	0.022150	0.000636	0.282567	0.000012	9.1	964	1092
KD-1.6	92	89	1.0	0.0643	0.0012	1.0952	0.0207	0.1235	0.0008	751	10	751	5	99%	0.024133	0.000806	0.282547	0.000012	8.3	996	1142
KD-1.7	91	102	0.9	0.0668	0.0013	1.1330	0.0223	0.1230	0.0008	769	11	748	5	97%	0.030109	0.000928	0.282509	0.000012	6.9	1053	1231
KD-1.8	82	93	0.9	0.0653	0.0015	1.1227	0.0248	0.1252	0.0009	764	12	760	5	99%	0.036699	0.001126	0.282452	0.000012	4.7	1139	1366
KD-1.9	106	103	1.0	0.0650	0.0015	1.1245	0.0254	0.1260	0.0009	765	12	765	5	99%	0.032186	0.000933	0.282417	0.000014	3.6	1181	1437
KD-1.10	176	139	1.3	0.0649	0.0013	1.1321	0.0224	0.1266	0.0009	769	11	769	5	99%	0.030557	0.000893	0.282456	0.000013	5.0	1126	1348
KD-1.11	121	130	0.9	0.0641	0.0011	1.0993	0.0175	0.1244	0.0007	753	8	756	4	99%	0.020004	0.000556	0.282582	0.000012	9.6	941	1055
KD-1.12	106	114	0.9	0.0648	0.0014	1.1023	0.0232	0.1239	0.0008	754	11	753	5	99%	0.022199	0.000769	0.282532	0.000012	7.8	1016	1174
KD-1.13	88	107	0.8	0.0647	0.0014	1.1130	0.0233	0.1248	0.0008	760	11	758	5	99%	0.023217	0.000634	0.282572	0.000012	9.2	956	1079
KD-1.14	42.3	48.7	0.87	0.0678	0.0025	1.1549	0.0407	0.1245	0.0015	780	19	757	8	97%	0.019023	0.000726	0.282529	0.000012	7.7	1020	1180
KD-1.15	186	174	1.07	0.0632	0.0014	1.0807	0.0250	0.1237	0.0009	744	12	752	5	98%	0.015930	0.000655	0.282552	0.000014	8.5	986	1127
KD-1.16	132	147	0.90	0.0668	0.0016	1.1394	0.0272	0.1231	0.0009	772	13	748	5	96%	0.020649	0.000778	0.282578	0.000013	9.4	952	1071
KD-1.17	557	319	1.75	0.0664	0.0013	1.1464	0.0228	0.1247	0.0011	776	11	757	6	97%	0.040835	0.001548	0.282564	0.000015	8.5	992	1128
KD-1.18	57.1	69.9	0.82	0.0652	0.0024	1.1234	0.0388	0.1250	0.0013	765	19	759	8	99%	0.018657	0.000706	0.282547	0.000013	8.3	993	1138
KD-1.19	96.4	115	0.84	0.0658	0.0016	1.1217	0.0259	0.1233	0.0011	764	12	750	6	98%	0.021640	0.000805	0.282541	0.000013	8.0	1005	1156
KD-1.20	80.6	73.7	1.09	0.0649	0.0023	1.1171	0.0377	0.1250	0.0014	762	18	759	8	99%	0.019432	0.000717	0.282551	0.000014	8.5	988	1129
KD-1.21	59.0	66.0	0.89	0.0662	0.0021	1.1389	0.0347	0.1246	0.0012	772	16	757	7	98%	0.020539	0.000765	0.282567	0.000013	9.0	967	1096
KD-1.22	168	170	0.99	0.0656	0.0016	1.1318	0.0267	0.1247	0.0010	769	13	757	6	98%	0.023090	0.000934	0.282542	0.000013	8.0	1006	1156

Fig. 3.

(a–b) Zircon CL images and concordia plots for the Shimian monzogranite and Kangding tonalite. Circles marked in zircons are analytical points with a diameter of 32 μm. (c) Zircon ε_Hf(t) vs. mean ²⁰⁶Pb/²³⁸U age plot. CHUR = chondrite uniform reservoir. Symbols and data sources as in Fig. 4. (d) Statistical histogram of ε_Hf(t) values.

Twenty zircons of sample SM-1 have ²⁰⁶Pb/²³⁸U ages of 846–811 Ma. Their weighted mean age is 827 ± 5 Ma (n = 20, MSWD = 4.2, 95% conf., Fig. 3a), and this is the best estimate for emplacement age of the Shimian monzogranite. Selected from sample KD-1, twenty-two zircons yield ²⁰⁶Pb/²³⁸U ages of 769–748 Ma. They have a weighted mean age of 757 ± 3 Ma (n = 22, MSWD = 1.18, 95% conf., Fig. 3b), which represents crystallization age of the Kangding tonalite.

Zircon Hf isotopes

Hf isotope data were calculated using weighted mean age of the investigated granitoids (Table 1). Zircons from sample SM-1 yield ε_Hf(t) values between 6.8 and –1.1, with two-stage Hf model ages (T_DM2) ranging from 1788 Ma to 1286 Ma. Zircons within sample KD-1 show ε_Hf(t) values of 9.6–3.6 and T_DM2 of 1437–1055 Ma (Figs. 3c and d).

Whole-rock geochemistry

Bulk rock chemical data are listed in Table 2. The Shimian monzogranite samples have high SiO₂ (76.05–70.87 wt.%), Al₂O₃ (14.26–12.74 wt.%), Na₂O (4.17–3.57 wt.%) and K₂O (4.55–1.80 wt.%) contents, with low Fe₂O₃^T (3.09–0.76 wt.%) and CaO (2.87–0.76 wt.%) concentrations. The Kangding tonalite samples show high SiO₂ (62.89–60.75 wt.%), Al₂O₃ (16.74–15.88 wt.%), Fe₂O₃^T (6.42–6.14 wt.%), MgO (2.52–2.33 wt.%), CaO (5.30–5.01 wt.%), Na₂O (3.97–3.66 wt.%) and K₂O (2.39–2.18 wt.%) contents. The rest major oxides (e.g., TiO₂, MnO and P₂O₅) of all samples are commonly lower than 1 wt.%. Furthermore, our samples are syenogranite and monzogranite illuminated by the CIPW normative mineral calculation (Fig. 4a). On the discriminant diagrams, all samples are granite to tonalite in composition, with calc-alkaline and metaluminous to peraluminous signatures (Figs. 4b–d).

Table 2. Bulk rock chemical contents of the Neoproterotoic Shimian monzogranite and Kangding tonalite

Samples	SM-1	SM-2	SM-3	SM-4	SM-5	SM-6	SM-7	SM-8	KD-1	KD-2	KD-3	KD-4	KD-5	KD-6	KD-7
Granitoids	The Shimian monzogranite								The Kangding tonalite
Major elements (wt.%)
SiO2	75.70	76.05	70.87	71.50	71.83	73.00	73.23	73.20	62.89	60.99	61.65	61.82	60.75	62.15	61.24
TiO2	0.11	0.08	0.39	0.40	0.31	0.31	0.26	0.25	0.75	0.68	0.71	0.71	0.68	0.72	0.69
Al2O3	12.74	13.00	14.26	14.06	13.88	13.88	13.78	13.85	15.88	16.74	16.30	16.50	16.60	16.55	16.62
Fe2O3T	1.44	0.76	3.06	3.09	2.56	2.61	2.30	2.11	6.14	6.42	6.24	6.26	6.19	6.31	6.21
MnO	0.02	0.02	0.07	0.06	0.06	0.05	0.05	0.06	0.12	0.11	0.12	0.12	0.12	0.12	0.12
MgO	0.12	0.13	0.89	0.79	0.65	0.59	0.54	0.57	2.33	2.52	2.41	2.44	2.48	2.47	2.47
CaO	0.76	1.18	2.79	2.87	2.21	2.33	1.85	2.08	5.01	5.30	5.18	5.22	5.30	5.24	5.27
Na2O	4.17	3.57	4.08	4.10	4.09	4.16	4.14	3.83	3.66	3.97	3.76	3.82	3.89	3.80	3.87
K2O	4.24	4.55	2.11	1.80	2.70	2.52	3.10	3.18	2.39	2.20	2.27	2.25	2.18	2.29	2.20
P2O5	0.02	0.01	0.11	0.11	0.07	0.08	0.06	0.06	0.17	0.20	0.18	0.18	0.19	0.18	0.19
LOI	0.41	0.47	1.00	1.10	0.83	0.95	0.76	0.75	0.58	0.72	0.80	0.74	0.79	0.72	0.77
Total	99.73	99.82	99.63	99.88	99.19	100.48	100.07	99.94	99.92	99.85	99.62	100.06	99.17	100.55	99.65
A/NK	1.11	1.20	1.58	1.62	1.44	1.45	1.36	1.42	1.84	1.88	1.89	1.89	1.89	1.90	1.90
A/CNK	0.99	1.00	1.01	1.01	1.01	1.00	1.02	1.02	0.90	0.90	0.90	0.91	0.90	0.91	0.91
FeOT/MgO	10.8	5.26	3.09	3.52	3.54	3.98	3.83	3.33	2.37	2.29	2.33	2.31	2.25	2.30	2.26
Na2O + K2O	8.41	8.12	6.19	5.90	6.79	6.68	7.24	7.01	6.05	6.17	6.03	6.07	6.07	6.09	6.07
K2O/Na2O	1.02	1.27	0.52	0.44	0.66	0.61	0.75	0.83	0.65	0.55	0.60	0.59	0.56	0.60	0.57
Mg#	14.2	25.3	36.6	33.7	33.5	31.0	31.8	34.9	43.0	43.8	43.4	43.6	44.3	43.7	44.1
Rare earth elements (ppm)
La	25.9	10.8	23.4	26.2	28.3	25.4	27.5	19.4	35.2	25.5	23.1	22.4	21.3	21.8	21.5
Ce	55.8	21.4	46.2	48.9	57.8	49.4	54.8	35.8	81.4	50.3	68.8	64.1	61.7	64.7	58.5
Pr	6.53	2.25	4.95	5.09	6.51	5.51	6.21	3.81	9.98	5.54	9.73	9.07	8.73	9.37	8.17
Nd	25.8	7.6	18.9	17.8	26.0	20.9	23.9	13.6	48.3	24.7	40.5	38.8	36.4	41.3	35.0
Sm	6.42	1.78	3.64	3.20	5.65	4.42	5.07	2.89	11.45	5.15	9.65	8.37	8.50	9.11	7.66
Eu	0.81	0.38	0.63	0.92	0.79	0.96	0.83	0.76	1.83	1.13	1.50	1.51	1.52	1.63	1.28
Gd	5.78	1.60	3.06	3.10	4.72	3.74	4.21	2.33	10.00	3.97	8.27	7.68	7.13	7.65	6.61
Tb	0.97	0.29	0.45	0.46	0.82	0.61	0.75	0.40	1.62	0.62	1.33	1.25	1.19	1.32	1.10
Dy	6.23	1.68	3.08	2.72	4.83	3.73	4.36	2.25	10.55	4.18	8.19	7.68	6.96	7.96	6.63
Ho	1.28	0.38	0.58	0.57	1.01	0.82	0.93	0.52	2.21	0.81	1.74	1.63	1.53	1.63	1.50
Er	3.60	1.09	1.84	1.70	2.74	2.26	2.53	1.40	7.09	2.48	5.13	4.65	4.50	5.04	4.20
Tm	0.52	0.17	0.24	0.26	0.41	0.35	0.38	0.23	1.08	0.38	0.80	0.73	0.70	0.77	0.65
Yb	3.34	1.18	1.76	1.62	2.53	2.25	2.39	1.64	6.89	2.42	5.19	4.81	4.49	4.83	4.35
Lu	0.48	0.18	0.26	0.27	0.38	0.35	0.36	0.28	0.97	0.38	0.78	0.76	0.70	0.81	0.64
ΣREE	143	50.8	109	113	142	121	134	85.3	229	128	185	173	165	178	158
(La/Yb)N	5.24	6.18	8.98	10.9	7.56	7.63	7.78	7.99	3.45	7.12	3.01	3.15	3.21	3.05	3.34
YbN	13.5	4.76	7.10	6.53	10.2	9.07	9.64	6.61	27.8	9.76	20.9	19.4	18.1	19.5	17.5
Eu/Eu*	0.40	0.68	0.56	0.88	0.46	0.70	0.53	0.87	0.51	0.74	0.50	0.57	0.58	0.58	0.54
Trace elements (ppm)
Rb	87.2	78.1	53.3	38.8	75.9	52.9	70.4	70.2	61.8	74.3	66.9	69.4	70.7	69.8	70.9
Sr	78.4	139	251	278	144	219	176	198	367	438	418	421	428	421	437
Ba	725	705	1005	1010	858	951	921	897	854	649	797	763	747	783	745
Nb	5.60	6.70	6.60	6.50	6.00	6.10	6.40	5.90	13.8	5.40	10.5	9.70	9.30	10.3	8.60
Ta	0.20	0.90	0.40	0.40	0.20	0.40	0.40	0.60	1.60	0.50	1.00	1.10	1.10	1.20	0.90
Zr	149	68.0	188	229	167	210	189	138	239	276	272	285	256	296	300
Hf	5.40	3.20	4.40	5.40	5.10	5.80	5.40	4.10	6.50	7.20	6.70	7.90	6.80	8.10	8.20
Th	7.73	17.6	5.78	5.68	7.74	6.54	7.33	12.2	5.17	3.88	5.03	5.06	4.89	4.72	4.60
U	1.17	6.06	1.02	1.57	1.12	1.50	1.10	2.89	1.70	1.38	1.44	1.41	1.47	1.41	1.44
Y	34.6	11.8	17.1	16.5	27.9	21.9	24.7	14.0	65.8	24.3	48.4	44.3	42.2	47.5	39.4
Ga	18.5	12.5	14.9	16.1	17.4	16.9	17.0	14.4	17.7	17.9	17.6	18.0	18.0	17.8	18.1
10000 × Ga/Al	2.74	1.82	1.97	2.16	2.37	2.30	2.33	1.96	2.11	2.02	2.04	2.06	2.05	2.03	2.06
Zr + Nb + Ce + Y	245	108	258	301	259	287	275	194	400	356	400	403	369	419	407
Y/Nb	6.18	1.76	2.59	2.54	4.65	3.59	3.86	2.37	4.77	4.50	4.61	4.57	4.54	4.61	4.58
Sr/Y	2.27	11.7	14.7	16.8	5.14	10.0	7.13	14.1	5.58	18.0	8.64	9.50	10.1	8.86	11.1
TZr (°C)	730	660	748	770	738	761	753	723	724	733	735	740	726	744	744

Fig. 4.

(a) Quartz-alkali feldspar-plagioclase (QAP) diagram. Abbreviation: 1 = quartzolite, 2 = alkali feldspar granite, 3a = syenogranite, 3b = monzogranite, 4 = granodiorite, 5 = tonalite, 6* = quartz-alkali feldspar syenite, 7* = quartz syenite, 8* = quartz monzonite, 9* = quartz monzodiorite/quartz monzogabbro, 10* = quartz diorite/quartz gabbro/quartz anorthosite, 6 = alkali feldspar syenite, 7 = syenite, 8 = monzonite, 9 = monzodiorite/monzogabbro, 10 = diorite/gabbro/anorthosite (Le Maitre, 1989). (b) FeO^T-(Na₂O + K₂O)-MgO diagram (Irvine and Baragar, 1971). (c) Total alkali vs. silica (TAS) diagram (Middlemost, 1994). (d) Molar Al₂O₃/(CaO + Na₂O + K₂O) vs. Al₂O₃/(Na₂O + K₂O) diagram (Maniar and Piccoli, 1989). (e–f) Normalized REE pattern and trace element diagram. Chondrite and primitive mantle values after Sun and McDonough (1989).

In general, REE and trace element patterns of the Shimian monzogranite resemble those of the Kangding tonalite (Figs. 4e and f). They all show high abundances of La, Ce and Nd, with various rare earth element contents (ΣREE) ranging from 229 ppm to 50.8 ppm. These granitoids are enriched in LREE with (La/Yb)_N ratios of 10.9–3.01 and Eu/Eu* values of 0.88–0.40 (Fig. 4e). They yield high Rb, Sr, Ba and Zr concentrations, and have enrichment in large ion lithophile elements and depletion in high field strength elements (Fig. 4f).

Discussion

Petrogenetic type

In this study, all the rock samples have low loss on ignition (LOI) values (1.10–0.41 wt.%) than alteration criterion (1.33 wt.%, Polat and Hofmann, 2003). In addition, the Shimian monzogranite and the Kangding tonalite show fairly homogenous major and trace compositions (e.g., Fig. 4). These features indicate that their compositions were not evidently influenced by later alteration.

The studied granitoids lack peraluminous minerals (Fig. 2), such as garnet and muscovite, and their A/CNK ratios are ranged from 1.02 to 0.90 (Fig. 4d). These signatures are incomparable to typical S-type granites, which are featured by Al-rich minerals and high A/CNK values of >1.1 (Chappell, 1999). Their Ga/Al ratios (2.74–1.82, Fig. 5a), FeO^T/MgO ratios (10.8–2.25) and Zr + Nb + Ce + Y contents (average 312 ppm) are commonly low, and are thus not in favour of typical A-type granites (Whalen et al., 1987). Previous studies have indicated that magma temperatures can be well estimated by zircon saturation thermometry using whole-rock compositions (Watson and Harrison, 1983). In this paper, new thermometer proposed by Boehnke et al. (2013) was applied to calculate zircon saturation temperatures, and yielded a range of 770–660°C (Table 2), which is well below the formation temperature of A-type granites (>800°C, King et al., 1997, Fig. 5b). Last but not least, the rock samples lack pyroxene grains that should commonly developed in A-type granites (Fig. 2). Hence, these lines of evidence indicate that S-type and A-type granites can be excluded for the studied granitoids.

Fig. 5.

(a) Y vs. Ga/Al diagram (Whalen et al., 1987). (b) T_Zr vs. Eu/Eu* diagram. (c) Molar CaO/(MgO + FeO^T) vs. K₂O/Na₂O diagram (Altherr and Siebel, 2002). (d) Rb/Ba vs. Rb/Sr diagram (Patiňo Douce, 1999). Symbols and data sources as in Fig. 4.

The Shimian monzogranite and the Kangding tonalite display metaluminous to weakly peraluminous signatures (Fig. 4d), and have declining Al₂O₃ with increasing SiO₂ (Fig. 6b). Moreover, hornblende grains generally occur in I-type granites, and they are also present in the investigated granitoids (Fig. 2). These features are comparable to typical I-type granites (Chappell, 1999; Chappell et al., 2012). Further, I-type granites are typically formed by partial melting of igneous source rocks (Chappell, 1999), and this is further illustrated by several discriminant diagrams (Figs. 5c and d). In conjunction, these studied granitoids fit the I-type definition.

Fig. 6.

Harker variation diagrams for the Neoproterotoic granitoids. Symbols and data sources as in Fig. 4.

Granitoid petrogenesis

The origin of I-type granites is still unclear. For instance, Wyborn et al. (1987) attributed them to fractional crystallization from primary mafic magmas. Barbarin (1996) proposed that they are associated with crustal partial melting and fractional crystallization. Moreover, Castro et al. (1991) suggested a mixing of mantle-derived and crustal melts.

Our samples are enriched in large ion lithophile elements and depleted in high field strength elements (Fig. 4f), indicating a crust-derived origin (e.g., Dai et al., 2017). They commonly have Y/Nb ratios higher than 1.2 (6.18–1.76, Table 2), showing a superficial derivation (Eby, 1992). In addition, Rapp and Watson (1995) indicated that crustal melts have Mg^# values of <40, and mantle-derived melts show Mg^# values of >40. The Mg^# values of our rock samples are 44.3–14.2, implying a major crustal origin. Synthetically, mantle-derived materials contributed little to the formation of the Shimian monzogranite and the Kangding tonalite.

Lu-Hf isotope systematics are sensitive geochemical tracers to decipher magma source with distinct signatures: (1) magma from depleted mantle or juvenile crust corresponding to positive ε_Hf(t) values higher than chondrite uniform reservoir (CHUR, ε_Hf(t) = 0); (2) magma from ancient crust exhibiting negative ε_Hf(t) values below CHUR; and (3) a mixed source showing variable (both positive and negative) ε_Hf(t) values (Kemp et al., 2007; Yang et al., 2007; Belousova et al., 2010). In this study, zircon Hf isotope data yield ε_Hf(t) values from 6.8 to –1.1 (average 3.6) for the Shimian monzogranite (Figs. 3c and d), suggesting a derivation from partial melting of juvenile crust with minor involvement of ancient crust. Their T_DM2 ages of 1788–1286 Ma are much older than the crystallization age of 827 Ma, indicating the contribution of ancient crustal components. In contrast, zircons within the Kangding tonalite show higher ε_Hf(t) values of 9.6–3.6 (average 7.8) and younger T_DM2 ages of 1437–1055 Ma (Table 1). This is suggestive of a primary juvenile crustal origin with negligible ancient crustal contamination. These point of views can be further reflected by Rb/Ba vs. Rb/Sr diagram, which confirms that more clay-poor constituents (such as greywacke) were involved in the magma source of the Shimian monzogranite (Fig. 5d).

The rock samples have variable SiO₂ contents of 76.05–60.75 wt.%, with strong correlations for several elements (Fig. 6), indicating significant fractional crystallization. Concretely, negative correlations between TiO₂ + Fe₂O₃^T and SiO₂ suggest Fe-Ti oxides (e.g., ilmenite) fractionation are likely present (Fig. 6a). Decreasing Al₂O₃, MgO and CaO with increasing SiO₂ were possibly related to fractionation of plagioclase and clinopyroxene (Figs. 6b–d). Moreover, the samples have Eu/Eu* of 0.88–0.40, (La/Yb)_N of 10.9–3.01 and Sr enrichment (Table 2). These data indicate that garnet was a main residual mineral (Defant and Drummond 1990).

In summary, we suggest a petrogenetic model for the investigated Shimian monzogranite and the Kangding tonalite, which were mainly derived from crustal melts with subsequent fractional crystallization. Specifically, their magma source was largely composed of juvenile crust (ε_Hf(t) >0) with minor ancient crustal components (ε_Hf(t) <0).

Neoproterozoic subduction and continental growth

The Neoproterozoic tectonic regime of the WYB is controversial. Different geodynamic models have been proposed based on various geochronological, geochemical and isotopic data of mafic-felsic rocks in this region, mostly focusing on slab-arc and mantle plume models. Zhou et al. (2006) investigated ca. 748 Ma Xuelongbao adakitic complex and proposed that the Neoproterozoic magmatism was subduction-related. They further pointed out that the South China craton was placed at the margin of the Rodinia supercontinent. Zhao et al. (2017) studied ca. 800 Ma Shimian ultramafic-mafic body in the WYB, and argued that it was most likely an incomplete Neoproterozoic SSZ-type (supra-subduction zone-type) ophiolite sequence. These studies strengthen the viewpoint of subduction process that has been extensively accepted (e.g., Zhao et al., 2008b, 2018; Du et al., 2014; Lai et al., 2015; Zhu et al., 2019). In contrast, plume model derives evidences from OIB-like basalts, mafic dykes and A-type granites in this area (Li et al., 1999, 2006). However, robust geological evidences such as extensive basalts, radiating dikes and crustal uplift are lacking (Zhao and Zhou, 2007).

According to Rb-Hf-Ta tectonic discriminant diagram, the Neoproterozoic I-type granitoids fall in the volcanic arc granite field, indicating a subduction setting (Fig. 7a). They show variable SiO₂ contents and Mg^# values, similar to those of the experimental slab melts (Fig. 7b). Moreover, they display low Al₂O₃ + MgO + TiO₂ + Fe₂O₃^T and high CaO + Al₂O₃ contents (Figs. 7c and d), implying melting at low pressures (Patiňo Douce, 1999). Combined with long-term magmatism and its linear distribution (Fig. 1), we are in favour of a slab-arc model in previous studies (e.g., Zhou et al., 2006; Zhao et al., 2018). Synthetically, we advocate that the Neoproterozoic I-type granitoids were most likely linked with arc magmatism and slab melts.

Fig. 7.

(a) Rb-Hf-Ta tectonic discriminant diagram (Harris et al., 1986). (b) Mg^# vs. SiO₂ diagram (Rapp et al., 1999). (c) (Al₂O₃ + CaO + Na₂O + K₂O) vs. Al₂O₃/(CaO + Na₂O + K₂O) diagram (Patiňo Douce, 1999). (d) CaO/Al₂O₃ vs. (CaO + Al₂O₃) diagram (Patiňo Douce, 1999). Symbols and data sources as in Fig. 4.

Continental growth is commonly attributed to lateral arc accretion (Desrochers et al., 1993) and vertical magma underplating (Peacock et al., 1994; Campbell, 2007). Previous research confirmed that zircon Hf isotope studies are useful for identifying continental growth (e.g., Belousova et al., 2010). In this paper, abundant zircons within I-type granitoids have positive ε_Hf(t) values above CHUR and even close to depleted mantle curve (Fig. 3c). This is indicative of significant continental growth for the Neoproterozoic WYB (Zhao et al., 2008a, 2018, 2021; Lai et al., 2015), and is compatible with the extensive occurrence of 860–750 Ma felsic and several mafic plutons in this region (Zhou et al., 2002, 2006; Zhao and Zhou, 2007; Zhao et al., 2018, 2021). Previous studies supported a northwestern marginal position for the Yangtze and Cathaysia blocks, as a key part of the Rodinia supercontinent (e.g., Zhao et al., 2011; Cawood et al., 2013; Zhu et al., 2021). Therefore, we propose lateral continental crustal growth resulted from ocean subduction along the Rodinia supercontinent margin (Barr et al., 1999; Zhao et al., 2008a, 2021).

Regional comparison and geodynamics

In the present study, our investigated Shimian monzogranite and Kangding tonalite are roughly adjacent (Fig. 1b), and both show I-type affinities (Figs. 4d, 5 and 6b). However, they have different formation ages and geochemical signatures. For instance, ca. 827 Ma Shimian monzogranite samples exhibit higher SiO₂ (Fig. 4c), TiO₂, Fe₂O₃^T, Al₂O₃, MgO and CaO contents (Fig. 6), with higher A/CNK ratios (Fig. 4d) and lower Mg^# values (Fig. 7b). They show relatively evolved Hf isotopes with ε_Hf(t) values varying between 6.8 and –1.1, whereas the Kangding tonalite samples are more depleted with ε_Hf(t) values of 9.6–3.6 (Figs. 3c and d). In order to detect these temporal-compositional variations and relevant geodynamic mechanism, we further collect published data of typical I-type granitoids in the Neoproterozoic WYB (Fig. 1b). These rocks are all calc-alkaline I-type with marked fractional crystallization and magmatic arc features in composition (Figs. 4b–d, 5, 6 and 7). Their zircon Hf isotopes are seated between 2.0–0.9 Ga crust evolution lines, with mostly positive ε_Hf(t) values (Figs. 3c and d). This is suggestive of a derivation from juvenile crust with minor ancient crust (Zhao et al., 2018).

Comparatively, all of the regional I-type granitoids including our samples have a long formation time span between ca. 828 Ma and ca. 748 Ma (Fig. 8). Their major element contents, trace element concentrations and Hf isotopes are extremely variable (Figs. 3–7), compatible with typical I-type granites in previous studies (e.g., Yang et al., 2007; Clemens et al., 2011; Gao et al., 2016). However, these variations are markedly bounded by ca. 780 Ma (Figs. 3c and 8), with contrasting SiO₂ contents (average 71.78 wt.% vs. 63.79 wt.%), Y/Nb ratios (average 3.91 vs. 3.49), Rb/Sr ratios (average 1.52 vs. 0.13), Mg^# values (average 27.9 vs. 43.6) and ε_Hf(t) values (average 4.84 vs. 7.76). This indicates different participation of juvenile crust in their source region before and after ca. 780 Ma (Eby, 1992; Rapp and Watson, 1995; Patiňo Douce, 1999; Kemp et al., 2007). Zhao and Zhou (2008) proposed that the regional 820–746 Ma heat anomaly was related to slab breakoff caused by rollback and steepening of subducted oceanic plate. It is reasonably speculated that slab breakoff may induce large-scale asthenosphere mantle upwelling (Atherton and Ghani 2002; Zhao et al., 2018), which heated the thickened crust and produced voluminous magmas (Zhao and Zhou, 2008). Additionally, Zhu et al. (2019) systematically studied ca. 781.1 Ma Dalu I-type granodiorite and ca. 779.8 Ma Dalu I-type granite, and further constrained that oceanic slab breakoff occurred at ca. 780 Ma in the WYB. Consequently, we attribute geochemical and isotopic diversities to ca. 780 Ma oceanic slab breakoff for the regional Neoproterozoic I-type granitoids.

Fig. 8.

Comparison diagrams for the Neoproterotoic granitoids. Symbols and data sources as in Fig. 4.

Based on evidences above, our preferred genetic scenarios can be addressed as follows. The Neoproterozoic Yangtze block was located at a northwestern marginal position of the Rodinia supercontinent (Fig. 9a). The oceanic crust was initially subducted eastward beneath the WYB at ca. 860 Ma (Du et al., 2014). During ca. 860–780 Ma the upwelling of slab melts formed juvenile crust, and partial melting subsequently took place under low pressure condition (Figs. 7c and d). Magmas ascended and were significantly contaminated by ancient crustal components. They underwent fractional crystallization and were emplaced in the upper crust (Fig. 9b). Eventually, I-type granitoids show higher SiO₂ contents (average 71.78 wt.%, Fig. 8a), Y/Nb ratios (average 3.91, Fig. 8c) and Rb/Sr ratios (average 1.52, Fig. 8d) with lower Mg^# (average 27.9, Fig. 8b) and ε_Hf(t) values (average 4.84, Fig. 3c). This is exemplified by ca. 828 Ma Qinganlin granite (Chen et al., 2015), ca. 790 Ma Shimian monzogranite (Zhao et al., 2008b), ca. 781.1 Ma Dalu granodiorite (Zhu et al., 2019) and ca. 779.8 Ma Dalu granite (Zhu et al., 2019). Slab breakoff was caused by rollback and steepening of subducted oceanic plate at ca. 780 Ma (Zhu et al., 2019). Subsequently, abundant juvenile crust was formed accompanying with large-scale asthenosphere mantle upwelling. Uplifting magmas were accordingly dominated by juvenile mafic melts, and the contribution proportion of ancient crust was sharply decreased after ca. 780 Ma (Fig. 9c). Ultimately, I-type granitoids such as ca. 757 Ma Kangding tonalite, ca. 754 Ma Luding monzodiorite (Lai et al., 2015) and ca. 748 Ma Luding granodiorite (Lai et al., 2015), display lower SiO₂ contents (average 63.79 wt.%, Fig. 8a), Y/Nb ratios (average 3.49, Fig. 8c) and Rb/Sr ratios (average 0.13, Fig. 8d), and higher Mg^# (average 43.6, Fig. 8b) and ε_Hf(t) values (average 7.76, Fig. 3c).

Fig. 9.

Petrogenetic model for the Neoproterotoic I-type granitoids in the WYB. Fig. 9a modified after Cawood et al. (2013).

Conclusions

1. The Shimian monzogranite and the Kangding tonalite yield emplacement ages of 827 ± 5 Ma and 757 ± 3 Ma, with ε_Hf(t) values ranging from 6.8 to –1.1 and from 9.6 to 3.6, respectively. Petrographic features and geochemical data indicate their I-type affinity, and derivation from juvenile crust with minor input of ancient crustal components.

2. The Neoproterozoic granitoids in the WYB were most likely associated with arc magmatism and lateral continental growth along the Rodinia supercontinent margin.

3. Regional comparison indicates that ca. 780 Ma slab breakoff was possibly responsible for compositional variations of the Neoproterozoic I-type granitoids.

Acknowledgments

We thank anonymous referees for insightful comments. This project was funded by the National Natural Science Foundation of China (41902068) and Young Scholars Development Fund of Southwest Petroleum University (201499010083).

References

•  Altherr,  R. and  Siebel,  W. (2002) I-type plutonism in a continental back-arc setting: Miocene granitoids and monzonites from the central Aegean Sea, Greece. Contrib. Mineral. Petrol. 143, 397–415.
•  Atherton,  M. P. and  Ghani,  A. A. (2002) Slab breakoff: A model for Caledonian, Late Granite syn-collisional magmatism in the orthotectonic (metamorphic) zone of Scotland and Donegal, Ireland. Lithos 62, 65–85.
•  Barbarin,  B. (1996) Genesis of the two main types of peraluminous granitoids. Geology 24, 295–298.
•  Barr,  S. R.,  Temperley,  S. and  Tarney,  J. (1999) Lateral growth of the continental crust through deep level subduction-accretion: A re-evaluation of central Greek Rhodope. Lithos 46, 69–94.
•  Belousova,  E. A.,  Kostitsyn,  Y. A.,  Griffin,  W. L.,  Begg,  G. C.,  O'Reilly,  S. Y. and  Pearson,  N. J. (2010) The growth of the continental crust: Constraints from zircon Hf-isotope data. Lithos 119, 457–466.
•  Boehnke,  P.,  Watson,  E. B.,  Trail,  D.,  Harrison,  T. M. and  Schmitt,  A. K. (2013) Zircon saturation revisited. Chem. Geol. 351, 324–334.
• BGMR (Bureau of Geology and Mineral Resources) of Sichuan Province (1991) Regional geology of Sichuan province. Geological Publishing House, 730 pp. (in Chinese with English abstract).
•  Campbell,  I. H. (2007) Testing the plume theory. Chem. Geol. 241, 153–176.
•  Castro,  A.,  Moreno-Ventas,  I. and  de la Rosa,  J. D. (1991) H-type (hybrid) granitoids: A proposed revision of the granite-type classification and nomenclature. Earth Sci. Rev. 31, 237–253.
•  Cawood,  P. A.,  Wang,  Y.,  Xu,  Y. and  Zhao,  G. (2013) Locating South China in Rodinia and Gondwana: A fragment of greater India lithosphere? Geology 41, 903–906.
•  Chappell,  B. W. (1999) Aluminium saturation in I-and S-type granites and the characterization of fractionated haplogranites. Lithos 46, 535–551.
•  Chappell,  B. W.,  Bryant,  C. J. and  Wyborn,  D. (2012) Peraluminous I-type granites. Lithos 153, 142–153.
•  Chen,  Q.,  Sun,  M.,  Long,  X. and  Yuan,  C. (2015) Petrogenesis of Neoproterozoic adakitic tonalites and high-K granites in the eastern Songpan-Ganze Fold Belt and implications for the tectonic evolution of the western Yangtze Block. Precambrian Res. 270, 181–203.
•  Clemens,  J. D.,  Stevens,  G. and  Farina,  F. (2011) The enigmatic sources of I-type granites: The peritectic connexion. Lithos 126, 174–181.
•  Clemens,  J. D.,  Regmi,  K.,  Nicholls,  I. A.,  Weinberg,  R. and  Maas,  R. (2016) The Tynong pluton, its mafic synplutonic sheets and igneous microgranular enclaves: The nature of the mantle connection in I-type granitic magmas. Contrib. Mineral. Petrol. 171, 1–17.
•  Dai,  Y.,  Zhu,  Y.,  Li,  T.,  Zhang,  H.,  Tang,  G. and  Shen,  Z. (2017) A crustal source for ca. 165 Ma post-collisional granites related to mineralization in the Jianglang dome of the Songpan-Ganzi Orogen, eastern Tiebtan Plateau. Geochemistry 77, 573–586.
•  Defant,  M. J. and  Drummond,  M. S. (1990) Derivation of some modern arc magmas by melting of young subducted lithosphere. Nature 347, 662–665.
•  Desrochers,  J. P.,  Hubert,  C.,  Ludden,  J. N. and  Pilote,  P. (1993) Accretion of Archean oceanic plateau fragments in the Abitibi greenstone belt, Canada. Geology 21, 451–454.
•  Du,  L.,  Guo,  J.,  Nutman,  A. P.,  Wyman,  D.,  Geng,  Y.,  Yang,  C.,  Liu,  F.,  Ren,  L. and  Zhou,  X. (2014) Implications for Rodinia reconstructions for the initiation of Neoproterozoic subduction at ~860 Ma on the western margin of the Yangtze Block: Evidence from the Guandaoshan Pluton. Lithos 196–197, 67–82.
•  Eby,  G. N. (1992) Chemical subdivision of the A-type granitoids: Petrogenetic and tectonic implications. Geology 20, 641–644.
•  Gao,  P.,  Zheng,  Y.-F. and  Zhao,  Z.-F. (2016) Experimental melts from crustal rocks: A lithochemical constraint on granite petrogenesis. Lithos 266–267, 133–157.
•  Guo,  J.-L.,  Gao,  S.,  Wu,  Y.-B.,  Li,  M.,  Chen,  K.,  Hu,  Z.-C.,  Liang,  Z.-W.,  Liu,  Y.-S.,  Zhou,  L.,  Zong,  K.-Q.,  Zhang,  W. and  Chen,  H.-H. (2014) 3.45 Ga granitic gneisses from the Yangtze Craton, South China: Implications for Early Archean crustal growth. Precambrian Res. 242, 82–95.
•  Harris,  N. B. W.,  Pearce,  J. A. and  Tindle,  A. G. (1986) Geochemical characteristics of collision-zone magmatism. Geological Society London, Special Publications 19, 67–81.
• He, J. M., Chen, G. H., Yang, Z. L., Min, J. K. and Liu, Q. X. (1988) The Gray Gneiss in Kangdian. Chongqing Publishing House, 243 pp. (in Chinese).
•  Hoskin,  P. W. O. and  Schaltegger,  U. (2003) The composition of zircon and igneous and metamorphic petrogenesis. Rev. Mineral. Geochem. 53, 27–62.
•  Irvine,  T. N. and  Baragar,  W. R. A. (1971) A guide to the chemical classification of the common volcanic rocks. Can. J. Earth Sci. 8, 523–548.
•  Jackson,  S. E.,  Pearson,  N. J.,  Griffin,  W. L. and  Belousova,  E. A. (2004) The application of laser ablation-inductively coupled plasma-mass spectrometry to in situ U-Pb zircon geochronology. Chem. Geol. 211, 47–69.
•  Kemp,  A. I. S.,  Hawkesworth,  C. J.,  Foster,  G. L.,  Paterson,  B. A.,  Woodhead,  J. D.,  Hergt,  J. M.,  Gray,  C. M. and  Whitehouse,  M. J. (2007) Magmatic and crustal differentiation history of granitic rocks from Hf-O isotopes in zircon. Science 315, 980–983.
•  King,  P. L.,  White,  A. J. R.,  Chappell,  B. W. and  Allen,  C. M. (1997) Characterization and origin of aluminous A-type granites from the Lachlan Fold Belt, Southeastern Australia. J. Petrol. 38, 371–391.
•  Lai,  S.,  Qin,  J.,  Zhu,  R. and  Zhao,  S. (2015) Neoproterozoic quartz monzodiorite–granodiorite association from the Luding–Kangding area: Implications for the interpretation of an active continental margin along the Yangtze Block (South China Block). Precambrian Res. 267, 196–208.
• Le Maitre, R. W. (1989) A classification of igneous rocks and glossary of terms: Recommendations of the International Union of Geological Sciences Subcommission on the Systematics of Igneous Rocks. Blackwell Science Inc, 206 pp.
•  Li,  X.,  Li,  Z.-X.,  Zhou,  H.,  Liu,  Y. and  Kinny,  P. D. (2002) U-Pb zircon geochronology, geochemistry and Nd isotopic study of Neoproterozoic bimodal volcanic rocks in the Kangdian Rift of south China: Implications for the initial rifting of Rodinia. Precambrian Res. 113, 135–154.
•  Li,  X.-H.,  Li,  Z.-X.,  Sinclair,  J. A.,  Li,  W.-X. and  Carter,  G. (2006) Revisiting the “Yanbian Terrane”: Implications for Neoproterozoic tectonic evolution of the western Yangtze block, South China. Precambrian Res. 151, 14–30.
•  Li,  Z. X.,  Li,  X. H.,  Kinny,  P. D. and  Wang,  J. (1999) The breakup of Rodinia: Did it start with a mantle plume beneath South China? Earth Planet. Sci. Lett. 173, 171–181.
•  Li,  Z. X.,  Bogdanova,  S. V.,  Collins,  A. S.,  Davidson,  A.,  De Waele,  B.,  Ernst,  R. E.,  Fitzsimons,  I. C. W.,  Fuck,  R. A.,  Gladkochub,  D. P.,  Jacobs,  J.,  Karlstrom,  K. E.,  Lu,  S.,  Natapov,  L. M.,  Pease,  V.,  Pisarevsky,  S. A.,  Thrane,  K. and  Vernikovsky,  V. (2008) Assembly, configuration, and break-up history of Rodinia: A synthesis. Precambrian Res. 160, 179–210.
•  Liu,  Y.,  Gao,  S.,  Hu,  Z.,  Gao,  C.,  Zong,  K. and  Wang,  D. (2010) Continental and oceanic crust recycling-induced melt-peridotite interactions in the Trans-North China Orogen: U-Pb dating, Hf isotopes and trace elements in zircons from mantle xenoliths. J. Petrol. 51, 537–571.
•  Maniar,  P. D. and  Piccoli,  P. M. (1989) Tectonic discrimination of granitoids. Geol. Soc. Am. Bull. 101, 635–643.
•  Middlemost,  E. A. K. (1994) Naming materials in the magma/igneous rock system. Earth Sci. Rev. 37, 215–224.
•  Patiňo Douce,  A. E. (1999) What do experiments tell us about the relative contributions of crust and mantle to the origin of granitic magmas? Geol. Soc. Lon. Spec. Pub. 168, 55–75.
•  Peacock,  S. M.,  Rushmer,  T. and  Thompson,  A. B. (1994) Partial melting of subducting oceanic crust. Earth Planet. Sci. Lett. 121, 227–244.
•  Polat,  A. and  Hofmann,  A. W. (2003) Alteration and geochemical patterns in the 3.7–3.8 Ga Isua greenstone belt, West Greenland. Precambrian Res. 126, 197–218.
•  Rapp,  R. P. and  Watson,  E. B. (1995) Dehydration melting ofmetabasalt at 8-32 kbar: Implications for continental growth and crust-mantle recycling. J. Petrol. 36, 891–931.
•  Rapp,  R. P.,  Shimizu,  N.,  Norman,  M. D. and  Applegate,  G. S. (1999) Reaction between slab-derived melts and peridotite in the mantle wedge: Experimental constraints at 3.8 GPa. Chem. Geol. 160, 335–356.
•  Sláma,  J.,  Košler,  J.,  Condon,  D. J.,  Crowley,  J. L.,  Gerdes,  A.,  Hanchar,  J. M.,  Horstwood,  M. S. A.,  Morris,  G. A.,  Nasdala,  L.,  Norberg,  N.,  Schaltegger,  U.,  Schoene,  B.,  Tubrett,  M. N. and  Whitehouse,  M. J. (2008) Plešovice zircon: A new natural reference material for U-Pb and Hf isotopic microanalysis. Chem. Geol. 249, 1–35.
•  Sun,  S.-s. and  McDonough,  W. F. (1989) Chemical and isotopic systematics of oceanic basalts: Implications for mantle composition and processes. Geol. Soc. Lon. Spec. Pub. 42, 313–345.
•  Watson,  E. B. and  Harrison,  T. M. (1983) Zircon saturation revisited: Temperature and composition effects in a variety of crustal magma types. Earth Planet. Sci. Lett. 64, 295–304.
•  Weidendorfer,  D.,  Mattsson,  H. B. and  Ulmer,  P. (2014) Dynamics of magma mixing in partially crystallized magma chambers: Textural and petrological constraints from the basal complex of the austurhorn intrusion (SE Iceland). J. Petrol. 55, 1865–1903.
•  Whalen,  J. B.,  Currie,  K. L. and  Chappell,  B. W. (1987) A-type granites: Geochemical characteristics, discrimination and petrogenesis. Contrib. Mineral. Petr. 95, 407–419.
•  Wiedenbeck,  M.,  Allé,  P.,  Corfu,  F.,  Griffin,  W. L.,  Meier,  M.,  Oberli,  F.,  Von Quadt,  A.,  Roddick,  J. C. and  Spiegel,  W. (1995) Three natural zircon standards for U-Th-Pb, Lu-Hf, trace element and REE analyses. Geostandard. Newslett. 19, 1–23.
•  Woodhead,  J. D. and  Hergt,  J. M. (2005) A preliminary appraisal of seven natural zircon reference materials for in situ Hf isotope determination. Geostand. Geoanal. Res. 29, 183–195.
•  Wyborn,  D.,  Turner,  B. S. and  Chappell,  B. W. (1987) The Boggy Plain Supersuite: A distinctive belt of I-type igneous rocks of potential economic significance in the Lachlan Fold Belt. Aust. J. Earth Sci. 34, 21–43.
•  Yang,  J.-H.,  Wu,  F.-Y.,  Wilde,  S. A.,  Xie,  L.-W.,  Yang,  Y.-H. and  Liu,  X.-M. (2007) Tracing magma mixing in granite genesis: In situ U-Pb dating and Hf-isotope analysis of zircons. Contrib. Mineral. Petr. 153, 177–190.
•  Zhao,  J.-H. and  Zhou,  M.-F. (2007) Geochemistry of Neoproterozoic mafic intrusions in the Panzhihua district (Sichuan Province, SW China): Implications for subduction-related metasomatism in the upper mantle. Precambrian Res. 152, 27–47.
•  Zhao,  J.-H. and  Zhou,  M.-F. (2008) Neoproterozoic adakitic suite at the northwestern margin of the Yangtze Block, China: Evidence for partial melting of thickened lower crust and secular crustal evolution. Lithos 104, 231–248.
•  Zhao,  J.-H.,  Zhou,  M.-F.,  Yan,  D.-P.,  Yang,  Y.-H. and  Sun,  M. (2008a) Zircon Lu-Hf isotopic constraints on Neoproterozoic subduction-related crustal growth along the western margin of the Yangtze Block, South China. Precambrian Res. 163, 189–209.
•  Zhao,  J.-H.,  Zhou,  M.-F.,  Yan,  D.-P.,  Zheng,  J.-P. and  Li,  J.-W. (2011) Reappraisal of the ages of Neoproterozoic strata in South China: No connection with the grenvillian orogeny. Geology 39, 299–302.
•  Zhao,  J.-H.,  Asimow,  P. D.,  Zhou,  M. F.,  Zhang,  J.,  Yan,  D.-P. and  Zheng,  J.-P. (2017) An Andean-type arc system in Rodinia constrained by the Neoproterozoic Shimian ophiolite in South China. Precambrian Res. 296, 93–111.
•  Zhao,  J.-H.,  Li,  Q.-W.,  Liu,  H. and  Wang,  W. (2018) Neoproterozoic magmatism in the western and northern margins of the Yangtze Block (South China) controlled by slab subduction and subduction-transform-edge-propagator. Earth Sci. Rev. 187, 1–18.
•  Zhao,  J.-H.,  Nebel,  O., and  Johnson,  T. E. (2021) Formation and evolution of a Neoproterozoic continental magmatic arc. J. Petrol. 62, egab029 (https://doi.org/10.1093/petrology/egab029).
•  Zhao,  X.-F.,  Zhou,  M.-F.,  Li,  J.-W. and  Wu,  F.-Y. (2008b) Association of Neoproterozoic A- and I-type granites in South China: Implications for generation of A-type granites in a subduction-related environment. Chem. Geol. 257, 1–15.
•  Zhao,  Z.-F.,  Gao,  P. and  Zheng,  Y.-F. (2015) The source of Mesozoic granitoids in South China: Integrated geochemical constraints from the Taoshan batholith in the Nanling range. Chem. Geol. 395, 11–26.
•  Zhou,  M.-F.,  Yan,  D.-P.,  Kennedy,  A. K.,  Li,  Y., Ding, J. (2002) SHRIMP U-Pb zircon geochronological and geochemical evidence for Neoproterozoic arc-magmatism along the western margin of the Yangtze Block, South China. Earth Planet. Sci. Lett. 196, 51–67.
•  Zhou,  M.-F.,  Yan,  D.-P.,  Wang,  C.-L.,  Qi,  L. and  Kennedy,  A. (2006) Subduction-related origin of the 750 Ma Xuelongbao adakitic complex (Sichuan Province, China): Implications for the tectonic setting of the giant Neoproterozoic magmatic event in South China. Earth Planet. Sci. Lett. 248, 286–300.
•  Zhu,  Y.,  Lai,  S.,  Qin,  J.,  Zhu,  R.,  Zhang,  F. and  Zhang,  Z. (2019) Geochemistry and zircon U-Pb-Hf isotopes of the 780 Ma I-type granites in the western Yangtze Block: Petrogenesis and crustal evolution. Int. Geol. Rev. 61, 1222–1243.
•  Zhu,  G.,  Liu,  H.,  Zhang,  T.,  Chen,  W.,  Xiao,  J.,  Zhao,  K. and  Yan,  H. H. (2021) Internal versus external locations of the South China Craton within Rodinia during the Cryogenian: Provenance history of the Nanhua Basin. Geol. Soc. Am. Bull. 133, 559–579.