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Article

Formation of Late Paleoproterozoic Gaositai Hornblendite in Northern North China Craton: Evidence from Zircon U-Pb Isotopes and Amphibole Trace Elements

by
Taichang Zhu
1,2,
Zhiwei Wang
1,2,*,
Zhihui Wang
1,2,
Yuxin Sun
1,2,
Zhenyu Liu
1,2,
Yin Xu
1,2,
Jingwen Yu
1,2,
Hao Wei
3 and
Xiaolei Geng
4
1
Hebei Key Laboratory of Strategic Critical Mineral Resources, Hebei GEO University, Shijiazhuang 050031, China
2
College of Earth Sciences, Hebei GEO University, Shijiazhuang 050031, China
3
Geological Testing Center, Hebei GEO University, Shijiazhuang 050031, China
4
Hebei Institute of Geological Survey, Shijiazhuang 050000, China
*
Author to whom correspondence should be addressed.
Minerals 2022, 12(8), 1046; https://doi.org/10.3390/min12081046
Submission received: 15 July 2022 / Revised: 12 August 2022 / Accepted: 17 August 2022 / Published: 19 August 2022
(This article belongs to the Special Issue Isotopic Tracers of Mantle and Magma Evolution)
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Abstract

:
Paleoproterozoic tectonic evolution of the northern North China Craton has been a hot research topic. We firstly identified a 1.85 Ga hornblendite from the Gaositai mafic–ultramafic complex, in northern Hebei. Systematic studies of petrology, zircon U-Pb geochronology, and in situ mineral major and trace elements of hornblendite are the key to revealing the petrogenesis of the Paleoproterozoic ultramafic rock and the tectonic evolution of northern North China Craton. LA-ICP-MS zircon U-Pb dating suggests the Gaositai hornblendite formed at 1851 ± 44 Ma. The late Paleoproterozoic ultramafic rocks, together with coeval post-collisional granites, formed a bimodal igneous assemblage. Both hornblende and its equilibrium melt compositions show strongly fractionated HREE patterns, relative enrichments in LREEs and LILEs, and depletions in HREEs and HFSEs. The phlogopite-bearing hornblendite magma could have originated from a hydrous garnet-facies mantle source metasomatized by slab-derived silicate melt. Furthermore, the variations of major and trace elements in hornblende from core to rim also reveal the mineral fractional crystallization and magma recharge. Zircon trace elements, melt composition equilibrium with hornblendes, and the bimodal igneous assemblage suggest that the generation of the Gaositai Paleoproterozoic hornblendite was likely the product of post-collisional extension related to the collision between eastern and western North China blocks.

1. Introduction

The North China Craton (NCC), one of the oldest and largest Precambrian cratons in China, develops the Archean to Paleoproterozoic crystalline basement, and Meso- to Neoproterozoic and Paleozoic sedimentary cover (Figure 1a). The NCC is a key object for studying the Precambrian tectonic regime transformation, evolution of supercontinents, and continental crust [1,2,3,4,5,6,7,8,9,10,11,12,13,14,15,16,17]. However, there is still a significant amount of disagreement regarding the timing of cratonization and the Paleoproterozoic tectonic evolution. For instance, some scholars believe that the North China Craton’s basement was created by numerous blocks colliding at 2.7–2.5 Ga, with rifting taking place throughout the Paleoproterozoic, and the blocks colliding once more at 1.95–1.85 Ga [15,18,19,20]. Another theory holds that the Yinshan and Ordos blocks, and the Longgang and Langrim blocks, amalgamated at 1.95 Ga and 1.90 Ga, respectively, to form the Western and Eastern blocks, and the oceanic crust was long-term subducted beneath the two blocks, but there were different understandings of subduction polarity. The Western and Eastern blocks ultimately collided together along the Trans-North China Orogen (TNCO) in 1.85 Ga, which coincided with the formation time of other major cratons throughout the world, representing the assembly of the Columbia supercontinent [3,4,8,21,22,23,24,25].
Neoarchean–Paleoproterozoic geological records are well-developed in the northern Hebei region of the NCC (Figure 1b); the predecessors focused mainly on the formation and metamorphic ages of the Neoarchean metamorphic complex, and petrogenesis of ~2.5 and 1.8 Ga metavolcanic rocks and granitoids, revealing the Neoarchean subduction and Paleoproterozoic collision to extension processes [9,13,25,26,27,28,29,30]. Traditionally, it has been assumed that there are large areas of Paleoproterozoic mafic-ultramafic rocks in the northern Hebei region, such as the Gaositai mafic–ultramafic complex, Damiao anorthosite, etc. [31]. Numerous studies have concluded that the Damiao anorthosite generated at 1.74 Ga, which was the product of post-collisional extension [24]. While the Gaositai mafic–ultramafic complex was considered to be a 280 Ma Alaskan-type rock association (Figure 1c), associated with an extensional mechanism that has different interpretations of back-arc extension and intracontinental extension [32,33,34], some researchers have also proposed that the formation and mineralization age was 213 Ma [35].
Minerals 12 01046 g001 550
Figure 1. Tectonic subdivision of the North China Craton, modified from Zhao et al. and Santosh [21,23] (a); simplified geological map of the Gaositai area in Northern Hebei (b); and a detailed geological map of Gaositai mafic-ultramafic rocks, modified from Chen et al. [33] (c).
Figure 1. Tectonic subdivision of the North China Craton, modified from Zhao et al. and Santosh [21,23] (a); simplified geological map of the Gaositai area in Northern Hebei (b); and a detailed geological map of Gaositai mafic-ultramafic rocks, modified from Chen et al. [33] (c).
Minerals 12 01046 g001
Fortunately, we discovered a 1.85 Ga hornblendite (CD1) from the Gaositai mafic-ultramafic rocks for the first time (Figure 1c). Previous studies have shown that magmatic zircons in hornblendites could record the formation age, zircon trace elements can be used to determine the tectonic setting, and the in situ major and trace element compositions of hornblendes and their equilibrium magma can reveal the nature of the primary magma and magma evolution [32,36,37,38,39,40,41]. Therefore, systematic petrology, zircon U-Pb dating, and in situ major and trace element analyses of hornblendite are the keys to revealing the petrogenesis of the Paleoproterozoic ultramafic rock and the geodynamic mechanism of the northern NCC.

2. Geological Setting and Sampling

The Gaositai area in northern Chengde is situated at the northernmost Trans-North China Orogen (TNCO), in the northern part of NCC (Figure 1a), and has developed a large area of Archean supracrustal rocks, granitic gneisses, Mesoproterozoic mafic rocks, and Late Paleozoic mafic-ultramafic rocks. The supracrustal rocks are represented by the Dantazi Group and the Hongqiyingzi Group. The former group is mainly distributed south of the Chicheng-Damiao-Pingquan fault, and the rock association includes trondhjemite-tonalite-granodiorite (TTG), mafic granulite, amphibolite, garnet-biotite schist, and marble. The latest research has determined that the formation ages of the TTG were Neoarchean to Paleoproterozoic (2.6–2.4 Ga), and most rocks in this group have undergone 2.4, 1.9–1.8 Ga granulite-amphibolite facies metamorphism [25,27,28]. The Paleoproterozoic Hongqiyingzi Group is distributed north of the fault, overlaid on the Dantazi Group in an angular unconformity, and the rock assemblage is gneiss, quartzite, mica schist, and marble. Liu et al. [9] recognized Neoarchean (2546–2532 Ma) quartz dioritic-tonalitic-granodioritic gneisses and Paleoproterozoic (1870–1819 Ma) granodiorite-monzogranite-syenogranite intrusions from the Hongqiyingzi Group, and some Late Paleozoic granites have been also identified from it [26]. Additionally, a small amount of 1894–1878 Ma S-type garnet granites occurred in Lanqi Town, Longhua, which is located north of the research region [29]. The Mesoproterozoic magmatic rocks include the Damiao 1742–1739 Ma anorthosite, mangerite, and minor troctolite [24]. Late Paleozoic mafic-ultramafic rocks also formed in Gaositai, Hongshila, and Boluonuo regions, including the 393–381 Ma pyroxenites, amphibole pyroxenites, hornblendites, and gabbros, as well as the 297–280 Ma gabbros [32,33,34].
In this study, we conducted petrological studies on a hornblendite (CD1) from the western part of the Gaositai mafic–ultramafic complex (Figure 1c), which intruded into the Archean–Paleoproterozoic metamorphic and deformed supracrustal rocks. The hornblendite has a massive structure, inequigranular texture, and a mineral composition dominated by amphibole (90%), magnetite (6%), pyrite (2%), and phlogopite (2%) (Figure 2). Hornblende can be divided into two groups. One is coarse-grained with clinopyroxene, with magnetite inclusions in the cores and pyrite inclusions in the rims. The fine-grained hornblende crystals in the other group coexist with magnetites and pyrites (Figure 2c,d). Phlogopite appears between medium- to fine-grained hornblende grains (Figure 2c). According to these petrological characteristics, we infer that clinopyroxene firstly crystallized, then the core of coarse-grained hornblende and magnetite was generated, followed by the formation of the mantle to the rim of coarse-grained hornblende, medium-to fine-grained hornblende, magnetite, pyrite, and phlogopite.

3. Analytical Methods and Results

3.1. Analytical Methods

Zircon U-Pb dating and trace element analysis were determined by laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) at Hebei Key Laboratory of Strategic Critical Mineral Resources, Hebei GEO University. The instrument couples a quadrupole ICP-MS (iCAP RQ) and a 193-nm ArF Excimer laser (RESOlution-LR). Laser spot size was set to 29 μm, laser energy density was 3 J/cm2, and the repetition rate was 6 Hz. A zircon 91500 standard was utilized for external age calibration, and a zircon GJ–1 standard was used as a secondary standard to monitor the deviation of age calculation. Calibrations for zircon trace element concentrations were performed using NIST SRM610 as an external standard and Si as the internal standard. Detailed instrumental conditions and data acquisition procedures were described by Wang et al. [42]. Data reduction was done using ICPMSDataCal and Isoplot (v.3.0) tools [43,44]. Isotope ratio and age uncertainties were quoted at the 1 sigma level. Zircon U-Pb dating and trace element results were given in Table 1 and Table 2.
Major and trace element compositions for hornblende were determined by LA-ICP-MS at Hebei Key Laboratory of Strategic Critical Mineral Resources. Laser spot size was set to 61 μm, laser energy was 100 mJ, and the repetition rate was 4 Hz. Data reduction was carried out by the iolite 4 software, using NIST SRM 610 as the external standard and Si as the internal one. NIST SRM612 was analyzed as a secondary standard quality control [45]. Major and trace element compositions of hornblende are reported in Table 3.

3.2. Analytical Results

3.2.1. Zircon U-Pb Dating

Zircon grains in the Gaositai hornblendite exhibit distinct cathodoluminescence (CL) image properties, and the corresponding U-Pb dating results likewise reveal multi-stage ages (Figure 3 and Figure 4a, Table 1). Three zircon grains, which range from subhedral to euhedral, exhibit oscillatory zoning and high Th/U ratios of 0.31–0.55, indicating a magmatic origin. Three analytical spots yield the youngest upper intercept age of 1851 ± 44 Ma (MSWD = 1.4) (Figure 4a; the calculated error correlation value is 0.6), which can represent the Gaositai hornblendite’s formation age. Another group of zircons mostly develops core-rim structures (Figure 3), and their inner cores with Th/U ratios of 0.27–1.4 show oscillatory zoning, which is of magmatic origin. However, their rims are zoneless, which points to a metamorphic origin. Seven spots of magmatic cores give an upper intercept age of 2206 ± 10 Ma (MSWD = 1.01). In addition, there are seven subrounded to rounded zircon grains in the sample, which show weak oscillatory zoning and relatively high Th/U ratios of 0.28–0.49, further indicating that these magmatic zircons underwent later metamorphism. Eight analyses yield one concordant 207Pb/206Pb age (2399 ± 13 Ma, MSWD = 0.17, n = 6) and two older 207Pb/206Pb concordant ages (2584 ± 27 Ma, 2522 ± 17 Ma).
The chondrite-normalized zircon rare earth element (REE) patterns show light rare earth element (LREE) depletion, heavy rare earth element (HREE) enrichment, and positive Ce anomalies and negative Eu anomalies (Figure 4b), consistent with the features of most magmatic zircons. Given the high Th/U ratios of zircons, this indicates that metamorphism has not significantly affected the REE patterns of 2.4–2.6 Ga zircons. In addition, it is noted that 1.85 Ga magmatic zircons contain relatively lower contents of heavy REE than the older captured zircons (Figure 4b).

3.2.2. In Situ Major and Trace Elements of Hornblende

We performed LA-ICP-MS in situ major and trace element analyses for coarse-grained hornblende and two medium-to fine-grained hornblende particles (Figure 2c and Figure 5, Table 3). In the chondrite-normalized REE diagram (Figure 5a), 16 analyses of hornblendes show upward convex REE patterns; LREE and medium rare earth elements (MREEs), particularly Pr, Nd, and Sm, are more enriched than HREEs. These REE patterns resemble those of hornblendes in many cumulates [47]. The primitive mantle-normalized trace element spidergram shows that the large ion lithophile elements (LILEs, such as Rb, Ba, K, Sr) are enriched, while Th, U, and HFSEs (Nb, Ta, Zr, Hf, Ti) are depleted (Figure 5b).
In particular, the trace element composition of the melt in equilibrium with the hornblende from the hornblendite was calculated by applying Am/LD values experimentally determined for basaltic systems [36,37] (Table 4). Equilibrium melts show strong fractionated REE patterns, as well as strong fractionation of MREEs and HREEs, enrichment of LREEs, LILEs, U, Pb, and depletion of HREEs, Nb, Ta, Zr, Hf, and Ti (Figure 5c,d).

4. Discussion

4.1. Formation Age of the Gaositai Proterozoic Hornblendite

Early studies established the Proterozoic as the age of formation for the Gaositai ultramafic rocks in the northern NCC [31]. Recent zircon U-Pb dating studies, however, have proposed the formation ages to be in the Paleozoic and Early Mesozoic. For instance, based on the age of a gabbro dike (280 Ma), Chen et al. [33] hypothesized that the Gaositai mafic-ultramafic rocks were formed in the Early Permian, and Zhang et al. [32] determined the crystallization age of a gabbro in Gaositai to be 392 Ma through SHRIMP zircon dating. In addition, the LA-ICP-MS zircon dating for chromitite and pyroxenite indicates that the formation age was the Late Triassic (213 Ma) [35]. Several Proterozoic igneous rocks have been discovered close to the research area, including the 1894–1878 Ma Lanqi garnet granite, 1870–1819 Ma granodioritic-monzogranites-syenogranites gneisses [9,29], the Damiao anorthosite pluton (1742–1739 Ma) [24], Lanying anorthosite and quartz syenite, and Gubeikou K-feldspar granite (1726 ± 9 Ma, 1739 ± 43 Ma, 1712 ± 15 Ma, and 1692 ± 19 Ma) [49]. However, the question remains whether there are any Proterozoic ultramafic rocks in the Gaositai and in nearby areas.
Therefore, we conducted LA-ICP-MS zircon U-Pb dating for a hornblendite in the Gaositai ultramafic rocks. The youngest zircons in the hornblendite are euhedral to subhedral, and develop magmatic oscillatory zoning (Figure 3) and high Th/U ratios of 0.31 to 0.55, indicating that they are all products of magmatic crystallization. The dating result indicates that the youngest upper intercept age of 1851 ± 44 Ma (Figure 4a) represents the forming age of hornblendite, this age is much older than the 1.7 Ga monzodiorite that intruded into the Gaositai Paleoproterozoic hornblendite and the Dantazi Group [31], our unpublished data. This Paleoproterozoic hornblendite should be separated from the Gaositai mafic–ultramafic complex that has Paleozoic zircon U-Pb ages. Thus, our study provides evidence for the presence of late Paleoproterozoic ultramafic rocks in the Gaositai area of northern Hebei; these rocks, along with the Paleoproterozoic granites, may form a bimodal igneous association. Furthermore, several 1894–1808 Ma mafic intrusive rocks also developed in southern Gaositai and in the nearby Chicheng areas [26,28], confirming the occurrence of coeval mafic magmatism in the eastern part of northern North China Craton.
On the other hand, the hornblendite samples contain three groups of magmatic captured zircons with ages of 2206 ± 10 Ma, 2399 ± 13 Ma, and 2584 ± 27 Ma (Figure 4a). The majority of the captured zircons modified by late metamorphism still exhibit relatively obvious magmatic oscillatory zoning and REE patterns of magmatic zircons (Figure 3 and Figure 4), which give ages that can be used to represent the formation ages of magmatic captured zircons, except for one zircon grain developing a thick zoneless metamorphic rim (2488 Ma). The coeval three stages of magmatic activity outlined above have also been found close to the study region. For instance, Feng et al. [50] reported S-type granites composed of garnet-biotite granite, garnet granite, and gneissic biotite granite in the Longhua area, which were formed at 2180 ± 42 Ma. The mafic dyke swarms in the Shimen area of eastern Hebei formed at 2162 Ma in an intracontinental rift environment [51]. In addition, several 2600–2505 Ma tonalites, trondhjemites, and granodiorites, as well as 2454–2404 Ma monzogranites and migmatites, have also been identified from the Dantazi Group and the Hongqiyingzi Group [25,26,28].

4.2. Magma Source and Evolution

Euhedral to subhedral hornblende crystals with cumulate structure in the Gaositai Paleoproterozoic hornblendite indicate the magmatic cumulate origin. The clinopyroxene and coarse hornblende were the early-crystallizing minerals, because the fine-grained clinopyroxene inclusions formed in the core of the coarse hornblende crystal. The hornblende compositions show no large variations, and no other anhedral silicate minerals crystallized from the interstitial melt. Hence, we attempt to qualitatively estimate the composition of equilibrium magma using the trace element composition of hornblende [36,37,41].
It should be noted that both hornblende and its equilibrium melt compositions show relative enrichment in LREEs and LILEs (e.g., Rb, Ba, K, Pb), and depletion in HREEs and HSFEs (e.g., Nb, Ta, Zr, Hf, Ti), which is geochemically similar to those of the hornblendes and their host basaltic rocks in subduction zones (Figure 5 and Figure 6). These characteristics might have been attributed to an origin of partial melting of a slab-derived fluid/melt metasomatized mantle source (Figure 6) [37,41,52,53,54]. Further partial melting under garnet-facies conditions comes from the observation of strongly fractionated REE and HREE patterns with lower HREE abundances in hornblende and equilibrium melt, as well as the relatively lower HREE contents in primary magmatic zircons (Figure 4 and Figure 5) [55,56]. In addition, the extensive crystallization of hornblende indicates the primary hydrous magma source, and the presence of phlogopite also gives the most direct evidence for mantle metasomatism, which is consistent with the negative correlations of Rb/Sr and Ba/Rb in the equilibrium melt [57]. U/Th and Th/Nb for mafic rocks have generally been used to determine distinct slab-derived components metasomatized by the mantle source, because LILEs and U are easily mobilized in slab-derived fluid, while LREEs and Th are more incompatible in slab-derived melts [53,54,58]. These two values for the equilibrium melt exhibit a notable slab-derived sediment-melt signature (Figure 6) [53]. This, together with the generation of phlogopite, suggests that the mantle source was metasomatized by slab-derived K-rich silicate melts.
Before emplacement, the magma underwent multiple stages of mineral fractional crystallization. The relatively low Cr, Co, Ni contents are consistent with early fractional crystallization of olivine and clinopyroxene. The mentioned mineral inclusions, as well as core-mantle composition changes of coarse-grained hornblende, including an increase of Ti and a decrease of Mg, Al, Ca, V (Figure 7), are evidence that an early fractional crystallization of clinopyroxene + large-grained hornblende (core) occurred under the conditions of relatively high water contents [56]. Then, the dominant mantle-rim of coarse hornblende + fine-grained hornblende + magnetite + pyrite + phlogopite crystallized, which resulted in the dominant decrease of Fe in hornblende from the residual magma [37,41]. In addition, the comparatively elevated and consistent Mg, Al, Ca, V, Sr/Y, La/YbN [48], and Dy/Yb found in late-forming hornblende rims and crystals (Figure 7), could be explained by the addition of mantle-derived magma [37]. Additionally, there is no plagioclase in hornblendite, and no obvious Eu anomaly or change in the hornblende and the equilibrium melt, indicating that there is no obvious plagioclase fractional crystallization, because the high water content of the magma suppressed the nucleation and crystallization of plagioclase [59,60,61].

4.3. Tectonic Setting and Implications

The tectonic setting of the generation of the Gaositai hornblendite is still not clear, as a result of the paucity of research on the formation of the 1.85 Ga mafic to ultramafic magmatic rocks in the study area. Thus, we use zircon trace elements and melt composition equilibrium with hornblendes to reveal the geodynamic mechanism.
Firstly, the high U/Yb and low Y concentrations of the magmatic and captured zircons in hornblendite set them apart from zircons from the oceanic crust (Figure 8a) [39]. Moreover, the presence of a large number of 2.2, 2.4, and 2.6 Ga magmatic captured zircons provides further evidence that the formation of the magmatic zircons was associated with the continental crust evolution. In addition, magmatic zircon trace elements can be used to determine the tectonic setting where magma was generated. For instance, in the U/Yb vs.Nb/Yb diagram, the magmatic zircons from MORB and OIB dominantly fall into the mantle array, and the magmatic zircons from the continental arc and post-collision rocks have relatively larger U/Yb ratios (Figure 8b) [40]. The 1851 Ma magmatic zircons from the hornblendite plot into the overlapping zone of continental arc and post-collisional origin zircons (Figure 8b) [40].
Furthermore, hornblende and equilibrium melt compositions have shown that the hornblendite magma originated from a hydrous garnet-facies mantle source metasomatized by silicate melt. Low Lu/Hf and high Th/Yb ratios of the equilibrium melt compositions indicate the addition of continental margin sediments (Figure 8c) [53]. At the same time, they show high Zr/Y, Th/Yb, Nb/Yb, and low Zr, which are similar to the characteristics of continental arc magma (Figure 6a and Figure 8d) [52,62]. These pieces of evidence point to the melt metasomatism in the mantle source region related to the partial melting of subducted continental sediments in a continental arc environment [53,54,58]. While the aforementioned geochemical characteristics could occur in both arc mafic rocks and post-collisional mafic to ultramafic rocks, the amphibole-rich rocks were mostly produced at convergent plate margins or distributed along orogenic belts, which could have formed in the oceanic ridge subduction or post-collisional extension settings [63,64,65,66]. In general, ridge subduction produced voluminous magmatic rocks, including rock assemblages of adakite, OIB, and Nb-rich basalt [67], which are absent in the study area. On the other hand, the notably slab-derived melt metasomatism in the mantle source at least implies that the slab subduction happened before the collision to post-collisional extension. Therefore, we consider that the generation of the Gaositai Paleoproterozoic hornblendite was likely the product of the post-collisional extension mechanism, and dominantly inherited the pre-subduction mantle source.
The North China Craton was formed by the collision of eastern and western North China blocks at 1.85 Ga [4,21,22,23,24,68]. Prior to the collision process, the Paleoproterozoic subduction probably occurred along TNCO, as indicated by the presence of Paleoproterozoic Chicheng SSZ ophiolite and the Paleoproterozoic cold-subduction related eclogite facies metamorphism [13,30,69]. To the north of the study area, 1894–1878 Ma S-type garnet-bearing granites [29], the post-collisional 1853–1827 Ma granodiorite- monzogranite-syenogranite intrusions, and 1742–1739 Ma Damiao anorthosite-norite-mangerite-troctolite have been recognized [9,24]. The 1853–1808 Ma gabbroic diorites and granites in the Chicheng region to the west formed in a post-collisional setting [70]. Additionally, the granulite facies metamorphism occurred in the Khondalite Belt and TNCO during 1.9–1.8 Ga [4,22,68,71]. These magmatic and metamorphic events documented the continent–continent collision between the eastern North China Block and the western North China Block, and the related post-collisional extension [4,9,22,23,24,29,70]. The extension process led to the asthenosphere upwelling, which triggered the partial melting of the garnet facies mantle rocks metasomatized by the slab-derived melt in the early stage. The predominant fractional crystallization and accumulation of hornblende in the underplating mafic magma beneath the lower crust resulted in the generation of late Paleoproterozoic hornblendite.

5. Conclusions

(1)
The Gaositai hornblendite formed at 1851 ± 44 Ma, together with coeval post-collisional granites, formed a bimodal igneous assemblage.
(2)
The magma of the Paleoproterozoic phlogopite-bearing hornblendite could originate from a hydrous garnet-facies mantle source metasomatized by slab-derived silicate melt, and has undergone mineral fractional crystallization and magma recharge.
(3)
The Gaositai Paleoproterozoic hornblendite formed in the post-collisional extension setting related to the collision between the eastern and western North China blocks.

Author Contributions

Conceptualization, Z.W. (Zhiwei Wang); methodology, Z.W. (Zhiwei Wang); software, T.Z. and Y.S.; investigation, T.Z., Z.W. (Zhiwei Wang), J.Y. and X.G.; data curation, T.Z., Y.S., Z.L. and H.W.; writing—original draft preparation, T.Z., Z.W. (Zhiwei Wang), Z.W. (Zhihui Wang), Y.S., Z.L. and Y.X.; writing—review and editing, T.Z., Z.W. (Zhiwei Wang) and Z.W. (Zhihui Wang); visualization, Z.W. (Zhiwei Wang); supervision, Z.W. (Zhiwei Wang); project administration, Z.W. (Zhiwei Wang); funding acquisition, Z.W. (Zhiwei Wang), H.W. and X.G.. All authors have read and agreed to the published version of the manuscript.

Funding

This research was funded by the Science and Technology Project of the Hebei Education Department (Grant No. ZD2021015), the Science and Technology Innovation Team project of Hebei GEO University (Grant No. KJCXTD-2021-05), the Central Guidance on Local Science and Technology Development Fund (Grant No. 216Z4201G), the National Natural Science Foundation of China (Grant No. 41802053), the Natural Science Foundation of Hebei Province (grant No. D2020403019), and the Opening Foundation of Hebei Key Laboratory of Strategic Critical Mineral Resources (Grant No. HGU-SCRM2201).

Institutional Review Board Statement

Not applicable.

Informed Consent Statement

Not applicable.

Data Availability Statement

The data presented in this study are available in this paper.

Acknowledgments

We are grateful to the handling editor, and thank the two anonymous reviewers for their constructive comments. We also thank Lu Yin and Kai Xing for their help with the LA-ICP-MS analysis.

Conflicts of Interest

The authors declare no conflict of interest.

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Figure 2. Photomicrographs of the Gaositai hornblendite. (a) Massive hornblendite, plane polarized light, (b) clinopyroxene and magnetite inclusions in hornblende, cross polarized light, (c) coarse- and fine-grained hornblendes, locations of in situ major and trace elements analyses, cross polarized light, (d) Magnetite and pyrite in hornblendite, reflected light. Cpx—clinopyroxene, Hb—Hornblende, Mag—magnetite, Phl—phlogopite, Py—pyrite.
Figure 2. Photomicrographs of the Gaositai hornblendite. (a) Massive hornblendite, plane polarized light, (b) clinopyroxene and magnetite inclusions in hornblende, cross polarized light, (c) coarse- and fine-grained hornblendes, locations of in situ major and trace elements analyses, cross polarized light, (d) Magnetite and pyrite in hornblendite, reflected light. Cpx—clinopyroxene, Hb—Hornblende, Mag—magnetite, Phl—phlogopite, Py—pyrite.
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Figure 3. Zircon CL images of the Paleoproterozoic Gaositai hornblendite. Numbers in circles are analysis numbers.
Figure 3. Zircon CL images of the Paleoproterozoic Gaositai hornblendite. Numbers in circles are analysis numbers.
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Figure 4. Ziron concordia diagram (a) and REE patterns [46] (b) of the Paleoproterozoic Gaositai hornblendite.
Figure 4. Ziron concordia diagram (a) and REE patterns [46] (b) of the Paleoproterozoic Gaositai hornblendite.
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Figure 5. Chondrite-normalized REE patterns (a,c) [46] and primitive mantle-normalized trace element spidergrams (b,d) [48] of hornblendes from the hornblendite and their equilibrium melt.
Figure 5. Chondrite-normalized REE patterns (a,c) [46] and primitive mantle-normalized trace element spidergrams (b,d) [48] of hornblendes from the hornblendite and their equilibrium melt.
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Figure 6. Th/Yb vs. Nb/Yb (a) [52] and U/Th vs. Th/Nb (b) [53] of equilibrium melt of hornblendes from the Paleoproterozoic Gaositai hornblendite.
Figure 6. Th/Yb vs. Nb/Yb (a) [52] and U/Th vs. Th/Nb (b) [53] of equilibrium melt of hornblendes from the Paleoproterozoic Gaositai hornblendite.
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Figure 7. Variations of major and trace elements and ratios from core to rim (a–j) of three hornblende grains (three colors) from the Gaositai Paleoproterozoic hornblendite.
Figure 7. Variations of major and trace elements and ratios from core to rim (a–j) of three hornblende grains (three colors) from the Gaositai Paleoproterozoic hornblendite.
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Figure 8. U/Yb vs. Y [39] (a), U/Yb vs. Nb/Yb [40] (b) for zircons, Lu/Hf vs. Th/Yb [53] (c), and Zr/Y vs. Zr [62] (d) for equilibrium melt of hornblendes from the Paleoproterozoic Gaositai hornblendite.
Figure 8. U/Yb vs. Y [39] (a), U/Yb vs. Nb/Yb [40] (b) for zircons, Lu/Hf vs. Th/Yb [53] (c), and Zr/Y vs. Zr [62] (d) for equilibrium melt of hornblendes from the Paleoproterozoic Gaositai hornblendite.
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Table
Table 1. LA-ICP-MS zircon U-Pb dating results for the Gaositai hornblendite.
Table 1. LA-ICP-MS zircon U-Pb dating results for the Gaositai hornblendite.
Spot No.ThUTh/UIsotopic RatiosAge(Ma)
ppmppm207Pb/206Pb1σ207Pb/235U1σ206Pb/238U1σ207Pb/206Pb1σ207Pb/235U1σ206Pb/238U1σ
CD1—11734030.40.15060.00149.64950.12010.46360.0046235416240212245520
CD1—226480.60.11430.00175.31010.10530.33610.0052186928187017186825
CD1—3861650.50.14580.00167.45410.09980.37010.0047229819216712203022
CD1—464930.70.13340.00193.72320.10000.20190.0048214424157622118626
CD1—5691590.40.16310.00119.45800.14780.41970.0062248812238314225928
CD1—630790.40.15500.00149.72850.20880.45230.0076240215240920240634
CD1—713310.40.17270.002811.86220.26510.49710.0081258427259421260135
CD1—86210.30.15530.00289.65710.20310.45190.0067240530240319240430
CD1—9923040.30.11230.00103.90490.05200.25160.0028183916161511144714
CD1—1019420.40.16640.001811.25070.26530.48950.0107252217254422256846
CD1—1113360.40.15820.00229.60980.20850.43920.0070243724239820234731
CD1—123632591.40.13930.00157.93240.10860.41200.0042221819222312222419
CD1—1314300.50.15390.00289.51920.21110.44810.0092239131238920238741
CD1—1416590.30.14200.00197.76690.12020.39580.0070225424220414215032
CD1—152072131.00.14570.00137.69070.12150.38170.0047229515219614208422
CD1—1742690.60.13790.00157.80900.09910.41090.0049221119220912221922
CD1—1825670.40.11020.00205.47030.16370.35790.0070180334189626197233
CD1—19681440.50.13270.00137.91350.10670.43190.0037220017222112231417
CD1—20831560.50.13270.00137.79050.10020.42550.0041220017220712228518
Table
Table 2. LA-ICP-MS zircon trace element results (ppm) for the Gaositai hornblendite.
Table 2. LA-ICP-MS zircon trace element results (ppm) for the Gaositai hornblendite.
Spot No.YNbLaCePrNdSmEuGdTbDyHoErTmYbLuThU
CD1—12512.510.0321.60.130.450.880.174.891.4518.17.7936.79.7310521.3173403
CD1—21622.480.0314.60.020.701.260.563.831.1815.05.5525.96.3966.214.726.347.5
CD1—32861.250.0014.90.070.741.180.185.671.9525.98.9542.211.111221.686.3165
CD1—43271.430.2514.40.652.281.810.266.712.3929.410.749.011.711821.663.992.8
CD1—52251.240.0013.00.070.390.790.153.681.4418.57.1434.59.1498.618.868.7159
CD1—62440.700.0010.80.040.240.860.204.351.4919.97.9137.69.9110420.129.878.5
CD1—73031.770.0015.40.150.671.330.146.362.1824.99.8446.311.411721.412.831.5
CD1—82461.250.0112.00.090.591.140.165.321.6620.97.8335.18.8892.516.55.7120.7
CD1—91481.890.0918.70.160.841.220.324.411.1813.64.5120.95.1752.810.291.7304
CD1—102981.610.0114.60.060.651.280.186.011.8824.69.4242.811.411621.318.542.0
CD1—112682.000.0015.60.050.641.270.165.471.8521.88.6539.610.010418.512.836.3
CD1—1212782.700.0635.00.252.626.210.7834.39.9312043.317940.735865.1363259
CD1—132591.280.0013.00.070.481.380.185.221.8122.78.3937.29.7110018.314.429.7
CD1—143160.960.0213.30.060.481.120.165.101.9423.59.6846.113.014425.915.959.0
CD1—159701.770.0424.90.504.718.330.5629.38.3198.732.913631.828748.3207213
CD1—1719086.190.0016.50.080.761.840.248.663.1538.214.768.417.017031.842.268.6
CD1—184761.790.0116.70.090.591.200.263.500.9710.33.5916.24.0341.58.1725.267.3
CD1—191130.690.0012.80.070.621.040.104.131.5419.97.8136.79.2493.318.868.5144
CD1—202471.320.0014.50.060.811.430.136.882.4530.211.852.813.613725.283.4156
Table
Table 3. Compositions (ppm) of amphiboles in the Paleoproterozoic Gaositai hornblendite.
Table 3. Compositions (ppm) of amphiboles in the Paleoproterozoic Gaositai hornblendite.
Spot No.HB-1HB-2HB-3HB-4HB-5HB-6HB-7HB-8HB-9HB-10HB-11HB-12HB-13HB-14HB-15HB-16
Li0.550.730.480.450.540.570.581.160.600.570.621.930.660.672.050.60
Na15,96216,63516,08316,07516,31716,44416,62017,71416,55616,42316,13816,64315,92115,94017,52116,338
Mg67,68667,89967,03567,53466,95667,41467,95768,05767,45667,12967,35165,93768,30567,94366,84367,998
Al61,15561,86460,76558,58559,01658,83661,15161,17659,78661,68161,39262,84361,09561,31962,04163,535
P64.0365.4870.2470.2476.5779.3971.0577.39100.45102.11123.9457.3366.8762.5765.4876.74
K9777970695498835872188739249868685719414988210,93110,09010,078912310,281
Ca79,87979,43278,83078,62777,78078,57079,53278,08477,48779,11778,87478,01279,04679,18879,71380,572
Sc10110099.298.897.599.296.791.588.995.9104110100100117102
Ti6046608462376007609560055942593760776198614962256053606157086180
V468502469442440438452495440454458492474468530495
Cr31.831.723.08.025.311.728.618.910.717.921.326.839.038.729.731.0
Mn863856891899901902853882917906901876894917883890
Fe79,23079,12182,84681,72083,60783,69377,00989,89084,96085,24984,35884,05281,60882,70283,26983,315
Co55.853.957.157.357.958.454.057.458.858.159.756.756.656.156.957.7
Ni24.123.623.221.822.320.923.861.924.124.025.826.525.024.832.126.6
Cu0.550.850.510.471.220.610.82758.410.650.762.502.810.690.520.680.53
Zn54.355.955.457.755.858.556.965.058.058.359.164.057.055.477.556.8
Ga14.414.613.713.514.013.714.115.714.014.213.915.314.214.416.014.7
Rb3.734.033.673.032.982.903.393.963.103.684.046.674.113.874.554.08
Sr387386377348353349389374360386386396377378374402
Y15.515.115.215.316.115.215.317.318.117.416.014.814.314.514.415.4
Zr14.814.614.813.314.014.315.918.318.517.916.115.515.015.816.516.0
Nb1.971.921.901.681.761.751.872.101.861.991.981.911.982.011.992.04
Sn0.820.930.860.830.770.790.901.061.021.040.970.880.880.951.151.04
Ba175180188176171170182191165186194201198189215199
La4.274.074.013.663.873.864.204.214.194.374.384.494.224.134.104.37
Ce14.514.414.113.614.414.214.515.416.216.215.715.714.514.414.415.2
Pr2.772.772.722.682.902.752.883.053.153.052.952.902.712.732.752.83
Nd16.015.514.715.616.215.716.117.218.217.316.515.815.315.415.515.8
Sm4.734.574.674.724.934.714.765.405.265.214.964.914.284.374.564.85
Eu1.481.441.461.561.581.501.471.591.701.681.571.531.401.351.381.46
Gd4.654.514.404.655.084.794.695.155.265.244.974.564.394.384.554.58
Tb0.620.640.610.610.640.630.630.700.720.680.640.620.570.610.580.64
Dy3.463.523.523.483.693.513.563.874.153.913.703.453.273.313.193.57
Ho0.660.610.640.650.670.650.640.720.780.700.660.630.570.600.590.64
Er1.641.701.661.621.641.611.591.851.991.741.681.581.601.441.381.63
Tm0.200.180.190.190.200.200.200.210.230.230.210.200.180.180.180.18
Yb1.091.041.101.111.301.211.141.421.491.311.071.090.991.021.011.08
Lu0.150.150.130.150.160.150.150.180.180.170.170.160.140.130.130.14
Hf1.121.071.091.031.101.041.261.391.441.391.231.151.101.121.191.09
Ta0.080.080.070.060.070.060.080.080.060.070.080.080.080.090.080.09
Pb0.720.800.760.720.770.760.771.560.740.810.810.870.730.760.750.85
Th0.060.060.060.050.050.050.050.050.050.060.060.070.070.060.070.06
U0.010.010.020.010.020.020.010.020.010.010.010.030.020.020.020.02
Table
Table 4. Melt compositions (ppm) in equilibrium with the amphiboles in the Paleoproterozoic Gaositai hornblendite.
Table 4. Melt compositions (ppm) in equilibrium with the amphiboles in the Paleoproterozoic Gaositai hornblendite.
Spot No.HB-1HB-2HB-3HB-4HB-5HB-6HB-7HB-8HB-9HB-10HB-11HB-12HB-13HB-14HB-15HB-16
Ti1652166317051642166616411624162316611694168117011654165615601689
K36,27837,80636,55236,53537,08337,37337,77340,25837,62737,32436,67637,82436,18536,22639,82137,132
Rb41.745.041.033.933.332.337.944.234.641.145.174.545.843.250.845.5
Sr622621608560568561627602579622622637607608602647
Y11.110.810.911.011.510.911.012.413.012.511.510.610.310.410.311.1
Zr32.932.332.829.431.031.735.240.541.139.635.734.333.335.036.735.4
Nb5.725.595.534.885.115.085.446.105.415.775.755.565.765.855.785.92
Ba475489512480465461495520449506527547537515586540
La23.722.622.320.321.521.523.323.423.324.324.425.023.523.022.824.3
Ce48.748.347.445.648.147.548.651.754.154.452.752.548.548.448.451.1
Pr6.026.025.915.836.315.976.256.636.856.626.426.315.895.925.986.16
Nd24.924.022.824.225.124.425.026.728.326.825.724.623.723.924.024.5
Sm4.484.334.434.484.674.464.515.114.984.944.704.654.054.144.324.60
Eu1.541.501.521.621.641.561.531.651.771.751.631.591.461.401.431.51
Gd3.533.423.343.533.863.643.563.913.993.983.773.463.333.323.463.48
Tb0.450.460.440.440.460.450.450.500.520.490.460.440.410.440.420.46
Dy2.432.482.472.442.592.472.502.722.912.752.602.422.302.322.242.51
Ho0.460.430.450.460.470.460.450.510.550.500.470.440.400.420.410.45
Er1.231.271.241.211.221.201.191.381.481.301.251.181.201.071.031.22
Yb0.930.890.940.951.111.030.981.221.281.130.920.940.850.880.870.92
Lu0.130.130.110.130.140.130.130.160.150.140.150.140.120.110.110.12
Hf1.481.411.441.361.451.381.661.841.901.831.621.511.451.471.571.44
Ta0.270.250.230.200.210.190.270.260.200.220.260.270.260.270.250.29
Pb6.016.736.396.056.426.416.4213.106.186.756.767.286.146.406.257.11
Th1.721.841.871.461.481.641.471.481.421.901.872.262.001.862.031.76
U0.350.480.620.510.630.540.430.550.520.440.471.030.650.630.630.64
Note: Am/LD values experimentally determined for basaltic systems are Ti 3.66, K 0.44, Rb 0.09, Sr 0.62, Y 1.39, Zr 0.45, Nb 0.34, Ba 0.37, La 0.18, Ce 0.30, Nd 0.64, Sm 1.06, Eu 0.96, Gd 1.32, Dy 1.42, Er 1.34, Yb 1.16, Lu 1.15, Hf 0.76, Ta 0.32, Pb 0.12, Th 0.03, U 0.03 [36,37], Pr 0.46, Tb 1.39, and Ho 1.42 (conjectured values of data fitting in this paper). Calculation formula is LC = AmC/ Am/LD.
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Zhu, T.; Wang, Z.; Wang, Z.; Sun, Y.; Liu, Z.; Xu, Y.; Yu, J.; Wei, H.; Geng, X. Formation of Late Paleoproterozoic Gaositai Hornblendite in Northern North China Craton: Evidence from Zircon U-Pb Isotopes and Amphibole Trace Elements. Minerals 2022, 12, 1046. https://doi.org/10.3390/min12081046

AMA Style

Zhu T, Wang Z, Wang Z, Sun Y, Liu Z, Xu Y, Yu J, Wei H, Geng X. Formation of Late Paleoproterozoic Gaositai Hornblendite in Northern North China Craton: Evidence from Zircon U-Pb Isotopes and Amphibole Trace Elements. Minerals. 2022; 12(8):1046. https://doi.org/10.3390/min12081046

Chicago/Turabian Style

Zhu, Taichang, Zhiwei Wang, Zhihui Wang, Yuxin Sun, Zhenyu Liu, Yin Xu, Jingwen Yu, Hao Wei, and Xiaolei Geng. 2022. "Formation of Late Paleoproterozoic Gaositai Hornblendite in Northern North China Craton: Evidence from Zircon U-Pb Isotopes and Amphibole Trace Elements" Minerals 12, no. 8: 1046. https://doi.org/10.3390/min12081046

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