The role of basaltic underplating in the evolution of the lower continental crust

Journal Pre-proofs The role of basaltic underplating in the evolution of the lower continental crust Jun Hu, Neng Jiang, Jinghui Guo, Wenbo Fan, Danqing Liu PII: DOI: Reference:

S0016-7037(20)30103-4 https://doi.org/10.1016/j.gca.2020.02.002 GCA 11641

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Geochimica et Cosmochimica Acta

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29 October 2019 29 January 2020 3 February 2020

Please cite this article as: Hu, J., Jiang, N., Guo, J., Fan, W., Liu, D., The role of basaltic underplating in the evolution of the lower continental crust, Geochimica et Cosmochimica Acta (2020), doi: https://doi.org/10.1016/ j.gca.2020.02.002

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The role of basaltic underplating in the evolution of the lower continental crust Jun Hua, b, c, Neng Jianga, b, c*, Jinghui Guoa, b, c, Wenbo Fanc, d, Danqing Liua, b, c aState

Key Laboratory of Lithospheric Evolution, Institute of Geology and

Geophysics, Chinese Academy of Sciences, Beijing 100029, China bUniversity cInnovation

of Chinese Academy of Sciences, Beijing 100049, China Academy for Earth Science, Chinese Academy of Sciences, Beijing

100029, China dState

Key Laboratory of Continental Dynamics, Northwest University, Xi’an, 710069,

China *Corresponding author: [email protected]

Abstract Basaltic underplating revealed by lower crustal xenoliths is usually considered to be manifested by younger zircon ages than those of the pre-existing crust and/or Sr–Nd isotopic heterogeneity resulting from mixing between mantle-derived basaltic melts and crust. The Hannuoba mafic lower crustal xenoliths have long been regarded as a typical example of Mesozoic underplating owing to the presence of 80–160 Ma zircons and evolved Sr–Nd isotopic compositions. However, our integrated study shows that most Mesozoic zircons in the Hannuoba mafic xenoliths precipitated from partial melts derived from the ancient lower crust. Although a few young zircons may record Mesozoic underplating, none of the xenoliths are products of Mesozoic underplating. Our study provides the first direct evidence that some zircons in lower

crustal xenoliths could be exotic. The contrasting O–Hf isotopic compositions of Mesozoic zircons from the Hannuoba lower crustal xenoliths allow us to distinguish zircons that were grown from different hydrous melts from those that represent recrystallized pre-existing zircons. This has major implications for geological interpretation of the age diversity commonly observed in deep-seated xenoliths worldwide. Furthermore, at Hannuoba both the ~1.8 Ga zircon ages from the granulite terrain and most of the 1.8–1.9 Ga zircon ages from a previously reported banded granulite xenolith also reflect metamorphism rather than underplating. It demonstrates that high-grade metamorphism or partial melting of Archean rocks can result in zircons with significantly younger U–Pb and THfDM ages. It is inferred that some granulite xenoliths previously regarded as products of Paleoproterozoic basaltic underplating from other regions (e.g., the Wyoming craton and the Siberian craton) may actually be remnants or derivatives of the pre-existing Archean lower crust. The large range in Sr–Nd isotopic compositions for various Hannuoba lower crustal xenoliths is unlikely to have resulted from mixing between basaltic melts and crust but was rather inherited from the ancient lower crust. Most of the Hannuoba mafic xenoliths can be best explained as residues left after partial melting of the late Archean lower crust that may be represented by the granulite terrain to produce the voluminous Mesozoic intermediate–felsic magmatic rocks. Therefore, young zircon ages and heterogeneous Sr–Nd isotopic compositions are insufficient criteria to infer recent underplating. Combined with literature data, it shows that Archean mafic granulite xenoliths are widespread globally, thus arguing against previous suggestion

that mafic granulites from the lower crust of most Archean cratons might have formed from post-Archean basaltic underplating. It is implied that the role of basaltic underplating in the evolution of the lower crust in many regions may need to be re-evaluated.

1.

INTRODUCTION

The lower crust plays an important role in the evolution of the continental crust. Granulite terrains and xenoliths provide two direct ways of investigating the nature of the relatively inaccessible lower crust. Globally, granulite terrains are dominantly of Archean ages and intermediate to silicic compositions with subordinate mafic compositions, whereas granulite xenoliths are mostly found in Mesozoic–Cenozoic basalts and are characterized by dominantly mafic compositions (Rudnick, 1992). Due to these differences, mafic granulite xenoliths from the lower crust of most Archean cratons are generally inferred to have formed from post-Archean basaltic underplating (Rudnick and Gao, 2014). If this is true, it raises the question whether the Archean mafic lower crust was never produced (hence different processes may have been operative in the formation of Archean crust), or it formed but was removed from the crust via density foundering (Kay and Kay, 1991). Abundant lower crustal and mantle xenoliths are found in alkali basalt components of the Hannuoba basalt. The basalt flows occupy >1700 km2 and were erupted between 22 and 10 Ma (Chen et al., 2001) through Archean granulite terrain in the northern North China craton (NCC) (Fig. 1). The age and compositional

differences are also evident between the granulite terrain and xenoliths in the Hannuoba region, leading to a suggestion that the Hannuoba mafic granulite xenoliths have formed by Mesozoic basaltic underplating (Fan et al., 1998; Zhou et al., 2002; Liu et al., 2004; Rudnick and Gao, 2014). However, some Hannuoba pyroxene-rich mafic granulite and Al-websterite xenoliths contain zircons with 1.8–1.9 Ga and 2.5–2.7 Ga ages (Zheng et al., 2004; Liu et al., 2004; Jiang and Guo, 2010), consistent with those of the granulite terrain and inherited zircons found in the Mesozoic magmatic rocks in the region. More importantly, the pyroxene-rich granulite xenoliths, the Al-websterite xenoliths and the Mesozoic magmatic rocks have almost identical Sr–Nd isotopic compositions, falling within the field of the granulite terrain. Obviously, any petrogenetic model for the Hannuoba mafic xenoliths must account for these similarities. 2.

GEOLOGICAL SETTING AND SAMPLES

The Hannuoba region is situated at the northern margin of the NCC (Fig. 1). Precambrian basement rocks (namely the Huai’an granulite terrain) are widely exposed in the region and consist mainly of granulites of mafic to silicic composition with subordinate amphibolite facies rocks. The rocks have been considered to represent an exposed lower crustal section (Zhai et al., 2001). U–Pb zircon dating yields both ~2.5 Ga and 1.8–1.9 Ga ages for rocks from the terrain. The ~2.5 Ga ages were considered to reflect their protolith age while the 1.8–1.9 Ga ages were interpreted to date metamorphism (Guo et al., 2005). Phanerozoic magmatism is widespread in the region and four episodes have been identified: (1) Devonian,

represented by the Shuiquangou syenitic complex (386–390 Ma) that covers an area of ca. 350 km2 and hosts several gold deposits (Miao et al., 2002; Jiang, 2005); (2) Triassic, represented by the Honghualiang granite (235 Ma, Jiang et al., 2007) and the Guzuizi granite (236 Ma, Miao et al., 2002); (3) Early Cretaceous, including intrusive and volcanic rocks with zircon ages of 125–143 Ma (Jiang et al., 2007, 2011); and (4) Tertiary, represented by the Hannuoba basalt flows. Xenoliths entrained in the Hannuoba basalts can be classified as peridotites, pyroxenites, and granulites (Wilde et al., 2003). Hannuoba granulite xenoliths include felsic, intermediate and mafic varieties, with mafic ones dominating. The mafic granulite xenoliths include garnet-bearing and garnet-free ones. The latter can be subdivided into pyroxene-rich and feldspar-rich (Both have similar SiO2, but the feldspar-rich xenoliths have lower Mg# [Mg/(Mg + ∑Fe)]) (Jiang and Guo, 2010). The pyroxene-rich granulite xenoliths have massive structure and granoblastic textures and consist dominantly of clinopyroxene (30-60%) and orthopyroxene (30-50%), with subordinate plagioclase (5-20%) (Fig. 2a). The plagioclases are often subhedral to anhedral and distributed between pyroxenes. The Hannuoba pyroxenite xenoliths have commonly been classified into an Al-augite group and a Cr-diopside group (Chen et al., 2001; Xu, 2002; Choi et al., 2008; Hu et al., 2016), following the classification of Wilshire and Shervais (1975). The Al-websterite xenoliths are dark black in hand specimen (Fig. 2b). They show similar massive structure and granoblastic textures as the pyroxene-rich granulite xenoliths. Meanwhile, they are also dominated by two pyroxenes with 40-65% clinopyroxene and 35-60%

orthopyroxene, but contain no plagioclase (Fig. 2c). A few Al-websterites may contain small amounts of tabular olivine (< 5%) and/or minor spinel (Fig. 2d). Both pyroxene-rich granulites and Al-websterites occur as discrete xenoliths. A total of 18 samples were investigated, including 8 from the granulite terrain (4 amphibolites and 4 mafic granulites), 6 pyroxene-rich granulite xenoliths and 4 Al-websterite xenoliths. The amphibolites are composed of amphibole, plagioclase, biotite and epidote, with minor quartz, garnet and opaques. Mafic granulites are composed of garnet, clinopyroxene, amphibole, plagioclase and minor opaques. The pyroxene-rich granulite and Al-websterite xenoliths are as described above. 3. ANALYTICAL METHODS 3.1. Major element analyses of minerals Major element compositions of the minerals were measured on polished thin-sections using a JEOL-JXA-8100 electron microprobe at the Institute of Geology and Geophysics, Chinese Academy of Sciences (IGGCAS). An energy-dispersive spectrometer was used for mineral phase identification while quantitative analyses were obtained by four wavelength-dispersive spectrometers. Several grains of each type of mineral were analyzed to obtain a representative average. The measurements were performed with an electron beam of 15 keV and 10 nA focused to 5 μm in diameter. 3.2. Major and trace element analysis of whole rocks In order to avoid potential contamination of the xenoliths by the host magma, the surrounding basaltic material was trimmed away from the xenolith with a diamond

saw. The remaining xenolith fragment was then washed and crushed to 40 mesh in a corundum jaw crusher. About 60 g of the sample split was powdered to less than 200 mesh using a vibratory tungsten carbide disc mill. The major element compositions of whole rocks were determined by X-ray fluorescence (XRF; Rikagu RIX 2100) using fused glass disks prepared from the sample powders at the Key Laboratory of Continental Dynamics, Northwest University, Xi'an. The trace element compositions of whole rocks were determined by quadrupole inductively coupled plasma mass spectrometry (Q-ICP-MS, Agilent 7500a) at the China University of Geosciences, Wuhan after acid digestion of samples in Teflon bombs. Analyses of standards and replicates are compiled in Table S3. 3.3. Sr–Nd–Hf isotope analyses of whole rocks The Sr and Nd isotopic compositions of most samples were determined on a Thermo-Fisher Triton TIMS at the Department of Terrestrial Magnetism (DTM), Carnegie Institution for Science. Strontium was measured using single Re filaments with Ta2O5–H3PO4–HF activator, while Nd was measured using double Re filaments. Both were measured using static multicollection on Faraday cups. Hafnium isotopic compositions were determined on a Nu-HR multicollector ICP-MS at DTM. Typical blanks were Sr = 80 pg, Nd = 60 pg and Hf = 70 pg, which were insignificant for the sample sizes used here. Measured signal sizes for the runs were 144Nd~3×10–11A. 146Nd/144Nd

Data are fractionation corrected to

86Sr/88Sr

88Sr~10–10A

and

= 0.1194 and

= 0.7219 using the exponential law. The average value for the NBS987 Sr

standard during the time period of these sample analyses was 87Sr/86Sr = 0.710263 ±

0.000004 (2 SD, n = 6) and for the JNdi Nd standard it was 143Nd/144Nd = 0.512111 ± 0.000002 (2 SD, n = 6). The Hf isotopic compositions are fractionation corrected to 179Hf/177Hf = 0.7325. Signal sizes for sample analyses ranged from 178Hf = 0.35 to 4.4 x 10–11A with all but four analyses having had

178Hf

signals above 1 x 10–11 A. The sample data were

collected in three analytical sessions over 4 months in 2010. Average values obtained for the JMC475 Hf standard run during the analytical sessions were 0.282153 ± 0.00006 (2 SD, n = 4, June), 0.282160 ± 0.000008 (2 SD, n = 2, July), and 0.282150 ± 0.000025 (2 SD, n = 3, September). The sample measurements were adjusted for the difference between the measured average value of the JMC-475 standard during each analytical session and the recommended value of 0.282160. Samples for which only Sr–Nd isotopic compositions are reported were measured using a Finnigan MAT 262 thermal ionization multicollector mass spectrometer at IGGCAS, following the procedures described by Li et al. (2016, 2019). Procedural blanks were <100 pg for Sm and Nd and <500 pg for Rb and Sr. The measured

143Nd/144Nd

and

87Sr/86Sr

ratios were normalized to

146Nd/144Nd

=

0.7219 and 86Sr/88Sr = 0.1194, respectively. The BCR-2 standard measured during the course of the analyses gave an average 143Nd/144Nd of 0.512626 ± 0.000013 (2 SD, n = 15), 147Sm/144Nd = 0.1382, Nd = 28.98 ppm, Sm = 6.61 ppm, 87Sr/86Sr = 0.705040 ± 0.000012 (2 SD, n = 6), 87Rb/86Sr = 0.4095, Rb = 46.30 ppm and Sr = 327.2 ppm. The samples were not leached prior to isotopic analysis. 3.4. Cathodoluminescence images, U–Pb–O–Hf isotope analysis of zircons

Cathodoluminescence (CL) images were obtained for zircons prior to age analysis, using a CAMECA SX-50 electron microprobe at IGGCAS. The O isotopic compositions of zircon were obtained using a CAMECA IMS-1280 at IGGCAS and are reported in the δ-notation relative to Standard Mean Ocean Water (SMOW). Detailed analytical procedures can be found in Li et al. (2010). The Cs+ primary beam was accelerated at 10 kV with an intensity of approximately 2 nA. The spot size was about 20 μm in diameter (10 μm beam diameter + 10 μm raster). An electron gun was used to compensate for sample charging during analysis. Secondary ions were extracted with a –10 kV potential. Oxygen isotopes were measured in multi-collector mode with two off-axis Faraday cups with each analysis consisting of 20 cycles × 4 s counting time. For zircon oxygen isotope analysis, the Penglai zircon with a δ18O value of 5.24 ± 0.26‰ served as the master reference standard to calibrate for analytical drift and instrumental mass fractionation. Point-to-point uncertainty was typically better than 0.4‰ (2 SD) for δ18O. During analyses, the average δ18O value of the standard zircon (Qinghu) was 5.42 ± 0.31‰ (2 SD, n = 16). The U–Pb dating of zircons was performed on the same spot as oxygen isotope analyses and conducted on a CAMECA IMS-1280 at IGGCAS. The spot diameter was 20–30 μm. Uncertainties in ages are quoted at the 95% confidence level. The analytical procedure has been described in detail previously in Li et al. (2009). The O2- primary beam was accelerated at –13 kV with an intensity of approximately 10 nA. In the secondary ion beam optics, a 60 eV energy window was used, together with a mass resolution of 5400. Each spot was measured over 7 scan cycles, and the total analytical time is ~12

min. Qinghu and Plesovice zircon standards are used for calibration and assessment of precision. The uncertainty in measured

206Pb/238U

ratios in a single analytical session

was generally around 1% (1 RSD) or less. Measured compositions were corrected for common Pb using non-radiogenic

204Pb

concentrations. Data processing was carried

out using Isoplot (Ludwig, 2012). The Hf isotopic analyses of zircons were conducted using the Neptune multi-collector ICP-MS, equipped with a Geolas 193 nm excimer laser, at the IGGCAS, Beijing. The analytical technique was described by Wu et al. (2006). Measurements of Lu-Hf isotope were obtained on the same grains that were previously analyzed for U-Pb and O isotopes, with ablation pits of 60 μm in diameter, ablation time of 26 s, repetition rate of 10 Hz, and laser beam energy density of 10 J/cm2. The εHf(t) and depleted mantle model ages (THfDM) are calculated assuming present day

176Lu/177Hf

and

176Hf/177Hf

ratios for average chondrite and the depleted

mantle to be 0.0336 and 0.282785 and 0.0384 and 0.28325, respectively (Blichert-Toft and Albarede, 1997; Bouvier et al., 2008). λLu = 1.867×10−11 year−1 (Söderlund et al., 2004). During the analyses, the

176Hf/177Hf

ratios of the standard

zircon GJ-1 and MUD were 0.282024 ± 0.000019 (2 SD, n = 8) and 0.282510 ± 0.000012 (2 SD, n = 8), respectively. 4. RESULTS The investigated 8 samples (including 4 amphibolites and 4 mafic granulites) from the granulite terrain have REE patterns and Sr–Nd isotopes comparable with previous results (Jiang et al., 2010) (Fig. 3a and Table S1). They yield εNd between –16.6 to –9.9. These 8 samples together with previously published 64 samples from

the granulite terrain define an isochron yielding an age of 2592 ± 72 Ma (MSWD = 0.26) (Fig. 4a), which is close to the ~2.5 Ga zircon ages reported from the granulite terrain (Zhao et al., 2005; Jiang et al., 2010). Most samples yield Nd depleted mantle model ages (TNdDM) of 2.5–2.8 Ga. They show a large range in 87Sr/86Sr and εNd (Fig. 4b). Five of the six investigated pyroxene-rich granulite xenoliths have convex upward REE patterns, consistent with previous results (Fig. 3b). However, the remaining sample JN0808 is distinct, with a LREE-enriched pattern similar to the granulite terrain rocks. In fact, a pyroxene-rich granulite xenolith 95DA6 with a similar LREE-enriched pattern but higher total REE contents (Fig. 3b) has been reported (Zhou et al., 2002). Nevertheless, they all have similar Sr–Nd isotopic compositions (Fig. 4b). The 4 Al-websterite xenoliths have 87Sr/86Sr of 0.706588–0.707061, εNd of –14.4 to –10.8 and convex upward REE patterns, consistent with literature data and almost identical to the pyroxene-rich granulite xenoliths (Figs. 3c and 4b). The zircon U-Pb ages and Hf–O isotopic compositions for Al-websterite xenolith HJ1435 are listed in Table 1. Ten of eleven zircon grains from Al-websterite HJ1435 yield Mesozoic ages and the remaining one gives a Precambrian 2163 Ma. The Mesozoic zircons have

206Pb/238U

207Pb/206Pb

age of

ages from 78 to 137 Ma, similar to

the dominantly Mesozoic (97–158 Ma) ages reported for an Al-websterite xenolith DMP-315 by Liu et al. (2004). Most of the Mesozoic zircons are concordant except that three are near concordant (spots 2, 8 and 9, Table 1, Fig. 5a). They can be divided

into two groups: Group I – 7 grains with low Th (2.2–109 ppm) and U (2.3–91 ppm) contents, ages of 78–132 Ma, and δ18O of 6.32–7.11‰; Group II – the remaining 3 with high Th (593–643 ppm) and U (573–719 ppm) concentrations, ages of 132–137 Ma, and δ18O of 5.22–5.48‰. Also, the two groups of zircons are significantly different in Hf isotopic compositions. Group I zircons have lower (0.00022–0.00059),

176Hf/177Hf

176Lu/177Hf

ratios

ratios (0.282032–282129) and εHf(t) values (–24.5 to

–21.0) than Group II zircons (176Lu/177Hf ratios = 0.00099–0.00148, 176Hf/177Hf ratios = 0.282887–282896 and εHf(t) values of 6.5–6.8). The Hf depleted mantle model ages (THfDM) of Group I zircons (1561–1691 Ma) are much older than those of Group II zircons (504–524Ma). The Precambrian zircon plots slightly below the Concordia and has a εHf(t) value of 1.3, a THfDM age of 2480 Ma and a δ18O value of 4.35‰. 5.

DISCUSSION

5.1. Hannuoba Al-websterite xenoliths: a lower crustal derivation The granulite xenoliths from Hannuoba are no doubt of lower crustal origin and the peridotite xenoliths are fragments of the upper mantle. The origin of the Al-websterite xenoliths, in contrast, is more enigmatic. Whereas a clear mantle derivation is evident for pyroxenites occurring as layers or dykes within peridotite massifs (Pearson and Nowell, 2004; Downes, 2007), the general lack of field relationships in the case of pyroxenite xenoliths makes it difficult to decide whether they are derived from the mantle or from the lower crust. In addition, owing to their bimineralic

(clinopyroxene

and

orthopyroxene)

assemblage,

no

suitable

geobarometers can be applied to confine their pressure conditions. Some pyroxenite

xenoliths are considered to be derived from the lower crust (Wysoczanski et al., 1995; Upton et al., 2001; Renna, 2008), whereas others are suggested to have a mantle derivation and result from metasomatism (Arai et al., 2006; Xu et al., 2008, 2010, 2013). Both a mantle derivation (Xu, 2002; Hu et al., 2016) and a lower crustal derivation (Liu et al., 2004; Wei et al., 2019) have been proposed for the Hannuoba Al-websterite xenoliths. For the mantle derivation, two different suggestions have been put forward: (1) High-pressure crystal-segregation from melts flowing through conduits in the mantle during the Mesozoic with the melts having been derived from a mixture of an asthenospheric melt and a delaminated continental crust (Xu, 2002); (2) Products of interaction between asthenosphere-derived silicate melt and the lithospheric mantle (Hu et al., 2016). For the lower crust derivation, the Al-websterites were considered to be products of Mesozoic basaltic underplating (Liu et al., 2004; Wei et al., 2019). Liu et al. (2004) suggested the xenoliths resulting from mixing between mantle-derived basaltic melts and pre-existing lower crust (i.e., the same origin as the Hannuoba mafic granulite xenoliths), whereas Wei et al. (2019) favored a derivation directly from an enriched mantle source. An enriched mantle derivation can be ruled out since there is no evidence of existence of an enriched mantle domain beneath Hannuoba. Our detailed comparisons show that the Al-websterite and pyroxene-rich granulite xenoliths are essentially indistinguishable in REE patterns (Figs. 3b and 3c), Sr–Nd isotopic compositions (Fig. 4b) and clinopyroxene mineral chemistry (not

shown, see Table S4). Their differences in both major and trace element concentrations, e.g., higher Al, Na, K, Rb, Sr, and Ba and lower Mg#, Ca, Cr and Ni for the pyroxene-rich granulite xenoliths (not shown, see Table S3), are due to the presence of plagioclase in the pyroxene-rich granulites. In addition, the Al-websterite and pyroxene-rich granulite xenoliths have similar age patterns (Fig. 6). These features suggest that the two types of xenoliths are cognate and both are components of the lower crust, thus arguing against a mantle derivation for the Hannuoba Al-websterite xenoliths. The Hannuoba pyroxene-rich granulite and Al-websterite xenoliths will be collectively referred to as mafic xenoliths hereafter unless otherwise specified. Below whether these mafic xenoliths are products of Mesozoic underplating will be examined in detail. It should be noted that the Hannuoba feldspar-rich mafic granulite xenoliths were not included in the above comparison. They have been considered to be cumulates of the Mesozoic granitoids (Jiang et al., 2009; Jiang and Guo, 2010) and have different origin from the pyroxene-rich mafic granulite xenoliths. 5.2. Origin of Mesozoic zircons in the Hannuoba lower crustal xenoliths Since Fan et al. (1998) published concordant

206Pb/238U

ages of 120–140 Ma for

two mafic and one intermediate granulite xenoliths from Hannuoba, subsequent works confirmed that both the Al-websterite and pyroxene-rich granulite xenoliths contain abundant Mesozoic zircons (Fig. 6) (Wilde et al., 2003; Liu et al., 2004; Jiang and Guo, 2010; Jiang et al., 2010, 2011; Wei et al., 2015), which are considered the strongest evidence for the Mesozoic underplating model (Fan et al., 1998; Chen et al.,

2001; Zhou et al., 2002; Liu et al., 2004; Rudnick and Gao, 2014; Wei et al., 2015, 2019). However, zircons with ~2.5–2.7 Ga ages but lacking Mesozoic ages were also reported for 3 mafic xenoliths JN0724, JN0813 and JN0737 (Jiang and Guo, 2010; Jiang et al., 2010, 2011). Interestingly, despite of different age patterns, the dated and undated mafic xenoliths have similar Sr–Nd isotopic compositions (Fig. 4b), pointing to a common origin. Clearly, defining the derivation of the Mesozoic zircons is critical for understanding the origin of the mafic xenoliths. The mafic xenoliths with ~2.5–2.7 Ga ages have zircon THfDM ages of 2.6–2.7 Ga and positive εHf(t) values close to that of the depleted mantle. Their whole rock εHf(130 Ma)

values plot near the 2.5 Ga average crust evolution line with

176Lu/177Hf

= 0.015

(Fig. 7), providing unambiguous evidence that they are remnants of late Archean lower crust. Thus, their whole rock εHf(130 Ma) values of –25.5 to –16.5 are likely to approximate that of the late Archean lower crust, which will be used as a reference to constrain the derivation of the Mesozoic zircons. Seven Mesozoic zircons from the Al-websterite HJ1435 have εHf(t) values of –24.5 to –21.0, essentially identical to those of Mesozoic zircons from 4 feldspar-rich mafic granulite xenoliths (–23.2 to –18.4, n = 60, Jiang and Guo, 2010; Jiang et al., 2011) and 4 felsic xenoliths (–23.3 to –19.1, n = 36, with an exception of –36.9, Wei et al., 2015) (Fig. 7, Table S8). The felsic and feldspar-rich mafic granulite xenoliths have been interpreted to be derivatives of the late Archean lower crust (Jiang and Guo, 2010; Jiang et al., 2011; Wei et al., 2019). Mesozoic zircons crystallizing from partial melts of the late Archean lower crust are expected to have εHf(t) values similar to the

putative εHf(130 Ma) of –25.5 to –16.5. Indeed, the Mesozoic zircons have εHf(t) values of –24.5 to –18.4, perfectly falling within the putative range and plotting tightly along the 2.5 Ga average crust evolution line (176Lu/177Hf = 0.015) (Fig. 7). This strongly indicates their crystallization from melts derived from the late Archean lower crust rather than from mantle-derived melts. Although no Hf isotope data on Mesozoic zircons have been reported for previously dated mafic xenoliths, the Al-websterite DMP0301B with exclusively Mesozoic zircons that yield an age of 158 ± 1 Ma (n = 5) has a whole rock εHf(130 Ma) value of –17.5 (Wei et al., 2015, 2019), falling within the inferred εHf(130

Ma)

range of –25.5 to –16.5 for the late Archean lower crust. This

demonstrates that the xenolith is most likely a remnant of the late Archean lower crust and its Mesozoic U–Pb zircon age does not necessarily indicate the xenolith resulting from recent underplating. However, the three Mesozoic (132–138 Ma) zircon grains with positive εHf(t) values (6.5–6.8) from the Al-websterite HJ1435 (Fig. 7) cannot be related to the late Archean lower crust. Instead, they could have crystallized from magmas derived from the depleted mantle. These zircons have δ18O values of 5.22–5.48‰, falling within the narrow range of δ18O (5.3 ± 0.3‰) for zircons in equilibrium with pristine mantle-derived melts (Valley, 2003). Therefore, these zircons may provide the first direct evidence of Mesozoic underplating at Hannuoba. However, they should be genetically unrelated to the xenolith. Apart from these few zircons, the xenolith shows no other evidence of input of depleted mantle components as the sample has whole rock Sr–Nd isotopic compositions virtually identical to those of the mafic xenoliths

with ~2.5–2.7 Ga zircon ages (Fig. 4b). Some Mesozoic zircons in the Hannuoba xenoliths could simply be reset late Archean grains. For example, zircons from an intermediate xenolith JN0919 (SiO2 of 60.33 wt.%) display core–rim structures (Jiang et al., 2011). The rims have concordant Mesozoic ages mostly between 120 and 145 Ma, approximating the ages of the Mesozoic intermediate–felsic rocks in the region, while the cores have discordant U–Pb ages with an upper intercept of 2501 ± 76 Ma, overlapping the zircon ages of the granulite terrain, and a lower intercept of 126 ± 13 Ma (Fig. 5b). The almost euhedral to subhedral shape of the zircons suggests that the rims were formed by recrystallization of pre-existing late Archean zircons. This is supported by the fact that both the rims and cores have identical THfDM ages of ~2.5 Ga (Table S8) and the Mesozoic zircons plot on the 2.5 Ga zircon evolution line (176Lu/177Hf = 0.001) (Fig. 7). In addition, the rims and cores have essentially identical δ18O values (6.58 ± 0.23‰ (n = 17) vs. 6.67 ± 0.33‰ (n = 15); Li and Jiang, 2012). Xenolith JN0919 has a whole rock TNdDM age of 2.45 Ga (Jiang et al., 2011). Therefore, xenolith JN0919 is no doubt a remnant of late Archean lower crust and its young zircon ages record metamorphism associated with the Mesozoic magmatism. In summary, Mesozoic zircons found in various Hannuoba lower crustal xenoliths have three different origins. They could have either crystallized from melts derived from the pre-existing lower crust, their late Archean ages may have been reset, or they may have crystallized from depleted mantle magmas. Although the Mesozoic zircons with positive εHf(t) from the Al-websterite HJ1435 may record Mesozoic underplating,

they are exotic to the xenolith. It is not clear how they were introduced into the xenolith. Nonetheless, they provide an example that some zircons can be genetically unrelated to the xenoliths. To our knowledge, this is probably the first direct evidence that young zircons in lower crustal xenoliths are exotic. The two groups of zircons in Al-websterite HJ1435 have similar Mesozoic ages but completely different εHf(t) and δ18O. None of them indicates that the xenolith is a product of Mesozoic underplating. The single zircon age of 2163 Ma from the Al-websterite HJ1435 yields a εHf value of -46.2 and a THfDM age of 2480 Ma. The zircon age is probably reset late Archean age. This may provide additional evidence that the xenolith has an ancient lower crustal derivation, rather than a product of Mesozoic underplating. 5.3. Additional evidence against a Mesozoic underplating model Since both the asthenospheric mantle from which the Hannuoba alkali basalts are considered to be derived (Song et al., 1990; Choi et al., 2008) and the lithospheric mantle that may be represented by the Hannuoba peridotite xenoliths (Rudnick et al., 2004) have depleted Sr–Nd isotopic compositions (Fig. 4b), the evolved Sr–Nd isotopic compositions of the Hannuoba mafic xenoliths are interpreted to result from mixing between mantle-derived basaltic melts and pre-existing lower crust by the Mesozoic underplating model (Liu et al., 2004). However, several lines of evidence argue against this mixing process. As a whole the various types of Hannuoba lower crustal xenoliths show a large range in Sr and Nd isotopic compositions (87Sr/86Sr = 0.70526 to 0.74624, εNd = –34.1 to 10.9), which is almost identical to that of the granulite terrain (Fig. 4b). In addition,

the 1.8–1.9 Ga and ~2.5 Ga zircon ages recorded in the mafic xenoliths (Fig. 6) are consistent with those of the granulite terrain (Zheng et al., 2004; Jiang and Guo, 2010; Jiang et al., 2010, 2011). Furthermore, except for a subduction and collision between the Eastern and Western blocks at 1.8–1.9 Ga (Zhao et al., 2005), the NCC was magmatically quiescent from ~2.5 Ga until a widespread tectonothermal reactivation in the Phanerozoic, indicating that most of the crust of the NCC was formed at ~2.5 Ga. In fact, the granulite terrain has been considered to represent an exposed late Archean lower crustal section (Zhai et al., 2001). If only rocks with SiO2<60 wt.% from the granulite terrain are considered, they have εNd mostly between –20 and –10, similar to the –19.4 to –10.8 values of the mafic xenoliths (Fig. 4a). Specifically, our investigated 8 samples from the granulite terrain have εNd of –16.6 to –9.9, even slightly higher than those of the mafic xenoliths. This strongly argues against the suggestion that the evolved Sr–Nd isotopic compositions of the Hannuoba lower crustal xenoliths might have resulted from mixing between mantle-derived basaltic melts and pre-existing lower crust (Zhou et al., 2002; Liu et al., 2004). Instead, they could be inherited from the pre-existing lower crust that may be represented by the granulite terrain. On the other hand, the various xenoliths do not fall on a mixing array on the

87Sr/86Sr

vs.

143Nd/144Nd

diagram (Fig. 4b) as expected for mixing

between mantle-derived basaltic melts and the pre-existing lower crust (Downes et al., 1990; Rudnick, 1990). Finally, unlike whole rock Sr–Nd isotopic compositions which are unable to retrieve the record of magmatic evolution, zircon has the advantage to trace changing

melt chemistry in its Hf and O isotope ratios (Griffin et al., 2002). The Mesozoic zircons interpreted to have crystallized from melts derived from the lower crust have a narrow range in both εHf(t) (Fig. 7) and δ18O (6.68 ± 0.24‰ for Al-websterite HJ1435 (Table 1), 6.77 ± 0.17‰ and 7.01 ± 0.13‰ for feldspar-rich mafic granulite xenoliths JN0831 and JN0815, respectively (Table S8)), providing robust evidence against mixing between depleted mantle-derived melts and pre-existing crust, otherwise a wide range of εHf and δ18O would be expected (Griffin et al., 2002). 5.4. A restite origin for the Hannuoba mafic xenoliths One may argue, that if the mafic xenoliths are derived from a lower crust as represented by the granulite terrain, they should plot on or close to the ~2.6 Ga isochron defined by the granulite terrain rocks in Fig. 4a. Indeed, two mafic xenoliths JN0808 and 95DA6 do plot on the ~2.6 Ga isochron. Whether these two xenoliths may just sample unexposed portions of the shallow granulite terrain should be addressed. Mineralogically, both xenoliths lack amphibole that is a ubiquitous mineral in mafic rocks from the granulite terrain. In addition, the two xenoliths have similar equilibration temperatures to other mafic xenoliths, higher than those of the terrain granulites (Chen et al., 2001; Guo et al., 2002). More importantly, xenolith 95DA6 contains zircons with both Precambrian and Phanerozoic ages (Wilde et al., 2003), with the latter being common in the mafic xenoliths but having never been reported from the granulite terrain. Given all these differences, xenoliths JN0808 and 95DA6 are unlikely derived from the shallow granulite terrain. Instead, they sample a considerably deeper portion of the crust, similar to other mafic xenoliths. They have

REE patterns and contents (Fig. 3b) as well as TNdDM ages (2886 Ma and 2522 Ma, respectively) similar to the granulite terrain rocks, indicating that the lower crustal portion from which they are derived is similar to the granulite terrain in both age and geochemistry. This provides additional evidence that the pre-existing lower crust could be represented by the granulite terrain. However, the remaining mafic xenoliths have higher

147Sm/144Nd

than the

granulite terrain rocks at given εNd and plot to the right of the ~2.6 Ga isochron (Fig. 4a). This can be best explained by partial melting of the ancient lower crust to produce the Mesozoic magmatic rocks in the region. During partial melting, Nd is expected to preferentially partition into the melt compared to Sm. As a result, the melts should have lower Sm/Nd ratios but higher incompatible element concentrations than the source rocks whilst the residues show the opposite. Both the melts and the residues should have similar (87Sr/86Sr)i and εNd(t) to the source rocks. All the expected features are exactly observed for the 125–143 Ma magmatic rocks (melts) and the mafic xenoliths (residues) if we assume partial melting of the source rocks geochemically similar to xenolith 95DA6 (Fig. 8). The assumption is reasonable because xenolith 95DA6 has both incompatible element concentrations and patterns comparable with the average of the granulite terrain rocks (Fig. 8). The significant difference in Ba between xenolith 95DA6 and the granulite terrain average is due to an unusually high Ba content (1300 ppm) in the xenolith. In fact, even higher Ba concentration (e.g., 1540 ppm for sample N10) can be found in granulite terrain rocks which have highly variable Ba (down to 25 ppm) (Table S1). Importantly, xenolith

95DA6 also has (87Sr/86Sr)i and εNd(t) (calculated at t = 130 Ma) consistent with the Mesozoic granitoids and the mafic xenoliths (Fig. 8). As shown in Fig. 8, both the Al-websterites and pyroxene-rich granulites have almost all incompatible elements lower than the lower crust values (Rudnick and Gao, 2014), consistent with a restite origin for the mafic xenoliths. Furthermore, the mafic xenoliths show (La/Yb)N negatively and TNdDM ages positively correlating with

147Sm/144Nd

(Fig. 9), indicating that they might have

undergone various degrees of partial melting. The higher the degree of the partial melting, the lower the LREE contents (thus higher

147Sm/144Nd)

in the residues.

Higher 147Sm/144Nd ratios would result in the unrealistically old TNdDM ages of 3.0–8.0 Ga for most of the mafic xenoliths, which reflect the modified nature of the late Archean lower crust. Xenolith JN0808 has a THfDM age of 2675 Ma similar to its TNdDM age and its whole rock εHf(130

Ma)

consistent with its

plots slightly above the 2.5 Ga average crust evolution line,

176Lu/177Hf

= 0.0178 (Fig. 7). These data show that xenolith

JN0808, along with xenolith 95DA6 which has a TNdDM age of 2522 Ma (THfDM age not available) (Zhou et al., 2002), may represent an unmodified portion of the late Archean lower crust. The whole rock εHf(130 Ma) of –19.5 for xenolith JN0808 further demonstrates the plausibility of the inferred εHf(130 Ma) range of –25.5 to –16.5 for the late Archean lower crust. Therefore, all studied mafic xenoliths, regardless of zircon ages, have a common late Archean lower crustal derivation on the basis of their similar Sr–Nd–Hf isotopic compositions (Figs. 4b and 7).

The Mesozoic igneous rocks from the Hannuoba region have magmatic zircon ages of 125–143 Ma, overlapping those of Mesozoic zircons from the mafic xenoliths (Fig. 6). Their magmatic zircon εHf(t) values mostly fall within the putative εHf(130 Ma) range of –25.5 to –16.5, indicating derivation dominantly from the late Archean lower crust. This suggests that the Mesozoic zircons with εHf(t) of ~–20 from the mafic xenoliths have the same origin as the magmatic zircons from the Mesozoic igneous rocks and both crystallized from melts derived from the late Archean lower crust. The Mesozoic magmatic rocks have inherited zircons ages (1.8–1.9 Ga and ~2.5 Ga) and εHf(t) values similar to those of the Paleoproterozoic and late Archean zircons from the mafic xenoliths and the granulite terrain (Figs. 6 and 7). The Mesozoic magmatic rocks have slightly higher 87Sr/86Sr than but similar εNd to the pyroxene-rich granulite and Al-websterite xenoliths (Fig. 4b). Given that the Mesozoic magmatic rocks have higher

87Rb/87Sr

ratios, when recalculated at t = 130 Ma, they will have (87Sr/86Sr)i

indistinguishable from the pyroxene-rich granulite and Al-websterite xenoliths. All these features suggest that the mafic xenoliths, the Mesozoic magmatic rocks and the granulite terrain are intrinsically linked. A restite origin for the Hannuoba mafic xenoliths, on one hand, demonstrates that the young zircon ages found in these xenoliths do not support a Mesozoic underplating model. It suggests that the role that Mesozoic underplating played in the evolution of the lower crust beneath Hannuoba may not be as significant as widely believed. On the other hand, it successfully explains the similarities in age and Sr–Nd–Hf isotopic compositions among the mafic xenoliths, the Mesozoic magmatic

rocks and the granulite terrain and the compositional complementarity between the Mesozoic magmatic rocks and the mafic xenoliths. While we eliminate the Hannuoba mafic xenoliths as products of Mesozoic underplating, an important issue remains unresolved. What caused the extensive Mesozoic partial melting of the ancient lower crust? Wilde et al. (2003) suggest that the heat needed for partial melting of the Archean lower crust may be provided by the extensive uprise of asthenospheric mantle. The three Mesozoic zircons with mantle-like Hf-O isotopic compositions found in Al-websterite xenolith HJ1435 could be taken as direct evidence of existence of Mesozoic underplating. The underplating may provide some heat for the partial melting. However, this evidence is only limited to zircon. We consider that the products of underplating should have depleted isotopic compositions. If mixing have taken place, its products should have more depleted isotopic compositions than the mafic xenoliths as the latter are the main components of the pre-existing lower crust. Further work is needed to find unambiguous xenolith evidence for the Mesozoic underplating. 5.5. Origin of 1.8–1.9 Ga zircons in the Hannuoba lower crustal xenoliths Zheng et al. (2004) have reported zircon U–Pb ages and Hf isotopic compositions for a granulite xenolith consisting of feldspar-rich (HNB3) and pyroxene-rich (HNB4) bands from Hannuoba. The banded xenolith yields zircon ages of 3.10 Ga (n = 1), 2.82 Ga (n = 1), 2.45–2.49 Ga (n = 2) and 1.73–1.96 Ga (n = 16). No younger ages were found. Most of the 1.8–1.9 Ga zircons have εHf(t) of –4.6 to 2.1 (Group I of HNB3 & 4, Fig. 7) and THfDM of 2.18–2.46 Ga. It is not straightforward to determine

whether these 1.8–1.9 Ga zircons reflect products of mixing between mantle-derived basaltic melts and the pre-existing lower crust during Paleoproterozoic underplating or merely a metamorphic event. However, it can be reconciled by comparison with studies of the granulite terrain previously conducted by Jiang et al. (2010). Both a mafic granulite JN0707 and an amphibolite JN0731 from the granulite terrain yield exclusively Paleoproterozoic zircon ages (mean 207Pb/206 ages of 1793 ± 13 Ma and 1814 ± 6 Ma, respectively). The zircons have εHf(t) values from –1.0 to 2.9 and –1.0 to 2.8 (Fig. 7) and THfDM of 2.08–2.21 Ga and 2.14–2.25 Ga, respectively. However, the two samples have whole rock TNdDM ages of 2.7–2.8 Ga. The decoupling between zircon THfDM (2.1–2.2 Ga) and whole rock TNdDM (2.7–2.8 Ga) has been interpreted to be caused by metamorphism (Jiang et al., 2010). As shown by Rubatto (2002) and Rubatto and Hermann (2003), crystallization or recrystallization of garnet during high-grade metamorphism can significantly deplete the associated medium in HREE to cause a HREE-depleted signature for co-precipitated minerals. Zircons grown under high-grade metamorphic conditions with garnet involvement are characterized by significantly declined Lu/Hf ratios but increased

176Hf/177Hf

ratios

(Zheng et al., 2005). Indeed, garnet is found in the two samples and the ~1.8 Ga zircons have remarkably low

176Lu/177Hf

ratios (0.000017–0.000054 for JN0707 and

0.000015–0.000037 for JN0731). Their declined 176Hf/177Hf

176Lu/177Hf

ratios and increased

ratios would result in enhanced εHf(t) values but decreased THfDM.

Therefore, both the zircon U–Pb and THfDM ages of mafic granulite JN0707 and amphibolite JN0731 cannot be used to infer their protolith ages. In fact, their whole

rock TNdDM of 2.7–2.8 Ga are essentially identical to that (2.72 Ga) of a tonalitic gneiss JN0703 from the granulite terrain which has a mean

207Pb/206

age of 2486 ±

20 Ma. Importantly, all of them plot on the ~2.6 Ga isochron defined by rocks from the granulite terrain (Fig. 4a). Therefore, the mafic granulite JN0707 and the amphibolite JN0731 from the terrain probably have a protolith age of 2.5–2.6 Ga and their ~1.8 Ga zircon ages reflect an event of metamorphism. Indeed, the ~1.8 Ga metamorphism has been widely reported from the basement of the NCC (Zhao et al., 2005). The results provide unambiguous evidence that high-grade metamorphism of Archean rocks can result in zircons with significantly younger U–Pb and THfDM ages. Compared with the ~1.8 Ga zircons from JN0707 and JN0731, the 1.8–1.9 Ga zircons from the banded granulite xenolith have slightly lower εHf(t) values (Fig. 7). Clearly, they also most likely reflect an event of metamorphism, rather than Paleoproterozoic underplating. The 2.45–2.49 Ga zircons from the banded granulite xenolith have εHf(t) of 5.7–10.1 and THfDM of 2.4–2.6 Ga, suggesting that the xenolith probably has a protolith age of ~2.5 Ga, similar to those of our mafic granulite xenoliths and the granulite terrain rocks. However, two zircons with ages of 1.89 Ga and 1.96 Ga from the pyroxene-rich band of the banded granulite xenolith have εHf(t) of 9.2 and 10.2 (Group II of HNB4, Fig. 7) and THfDM of 1.93 Ga and 1.96 Ga, respectively. They could have crystallized from magmas derived from the depleted mantle. Similar to the three Mesozoic zircons with mantle-like Hf–O isotopic compositions from the Al-websterite HJ1435, these

two zircons are considered to be exotic and genetically unrelated to the banded xenolith because both the remaining 1.8–1.9 Ga zircons and 2.45–2.49 Ga zircons point to a protolith age of ~2.5 Ga for the banded xenolith. 5.6. Application to other regions and implications Our proposed scenario can also be applied to the whole eastern NCC where widespread 120–140 Ma intermediate–felsic magmatic rocks commonly have ~2.5 Ga inherited zircons and their Sr–Nd isotopic compositions are similar to the Hannuoba counterparts (Jiang et al., 2011), suggesting derivation mainly from the Archean lower crust and existence of a common ancient lower crust. The intense magmatism may be a direct consequence of lithospheric loss beneath the eastern NCC in the Mesozoic (Menzies et al., 1993). A restitic mafic xenolith 83-159 from McBride, northern Queensland, Australia, has abundant concordant zircons with ages of 200–350 Ma and a number of discordant zircons that yield an upper intercept of 1520+150 -190 Ma and a lower intercept of 260+140 -100 Ma (Rudnick and Williams, 1987), similar to the age pattern of the Hannuoba intermediate xenolith JN0919 (Fig. 5b). Its (87Sr/86Sr)i and εNd(t) are similar to the ~300 Ma Yataga granodiorite in the region (0.712892 and –12.6 vs. ~0.713 and –12.5, t = 300 Ma) (Black and McCulloch, 1990; Rudnick, 1990). They could be an example of a residue–melt association. Given that the Mesozoic zircons from the Hannuoba lower crustal xenoliths have three different origins, the origin of the Paleozoic zircons in the McBride xenoliths might be tested by a combined study of zircon O and Hf as well as whole rock Sr–Nd–Hf isotopic

compositions. Also, some lower crustal xenoliths from the French Massif Central may also be genetically related to the Hercynian granitoids in that region. Lower-intercept ages of 280–300 Ma for zircons have been reported for some xenoliths which have (87Sr/86Sr)i and εNd(t) (t = 300 Ma) similar to the granitoids. Some meta-igneous xenoliths possess features of residues (Downes et al., 1990). Another possible example of correlation between lower crustal xenoliths and regional surface geology comes from the Camp Creek where similar Proterozoic Nd model ages between Camp Creek mafic xenoliths and nearby Cenozoic and Mesozoic granitoids led Esperança et al. (1988) to consider that the lower crust represented by the mafic xenoliths may be a suitable source for the younger plutonics. All these examples demonstrate that similar partial melting process of ancient lower crust must have occurred in all continental areas. Furthermore, the Hannuoba mafic xenoliths actually do have "modern" analogs as seen in the Sierra Nevada (Lee et al. 2006) and Kohistan arcs (Jagoutz et al., 2007). They appear remarkably similar to high MgO pyroxenites, interpreted to be deep level cumulates/restites associated with granitoid formation (Lee et al., 2006). These examples demonstrate the association between xenoliths and magmatism. Like the Hannuoba counterparts, lower crustal xenoliths from many other regions commonly have complicated age populations (Rudnick and Williams, 1987; Bolhar et al., 2007; Shatsky et al., 2019). Usually, the young zircon ages are interpreted to record basaltic underplating or related metamorphism (Rudnick and Williams, 1987).

However, we show that both the 1.8–1.9 Ga and 80–160 Ma zircons from the Hannuoba lower crustal xenoliths do not indicate that the xenoliths are products of either Paleoproterozoic or Mesozoic underplating. This finding is important because it could be erroneous to infer xenoliths to be products of recent basaltic underplating based on young zircon ages. Distinction between zircons of different origins has great bearing on geological interpretation of the age diversity commonly observed in deep-seated xenoliths. This finding has implications for the evolution of lower crust in other regions. Based on 1.75–1.81 Ga zircon ages from mafic granulite and garnet paragneiss xenoliths, Gorman et al. (2002) suggest that the lowermost 12 km of the lower crust beneath the Wyoming craton represent products of Paleoproterozoic underplating, thus significantly younger than the upper part of the crust which has a late Archean age. However, Bolhar et al. (2007) reported a lower crustal xenolith from Bearpaw Mountains in the northern Wyoming craton having mostly late Archean zircon ages with subordinate ~1.8 Ga ages. The ~1.8 Ga ages are interpreted to reflect Paleoproterozoic high-grade metamorphism. It is inferred that the 1.75–1.81 Ga zircon ages of the lower crustal xenoliths reported by Gorman (2002) might actually register metamorphism rather than underplating. Similarly, both Archean and Paleoproterozoic zircon ages were reported from lower crustal xenoliths in the Siberian craton (Moyen et al., 2017; Shatsky et al., 2016, 2019). In light of zircon U–Pb ages and Hf isotopic compositions of granulite xenoliths from Udachnaya of the Daldyn-Markha domain, Moyen et al. (2017)

conclude that the lower portions of the Archean lithosphere beneath the region were removed and replaced by Paleoproterozoic components and the upper and lower crust is decoupled. In fact, their 1.8–1.9 Ga zircons have a large range of εHf(t) values from –20.9 to –1.7, which can be divided into three groups (Group I: –20.9 to –16.1, Group II: –10.2 to –8.1 and Group III: –3.5 to –1.7). Group I zircons have the lowest εHf(t) and are interpreted to have recrystallized from older, late Archean grains under high-grade metamorphic conditions, the same origin as the Mesozoic zircons of the Hannuoba intermediate xenolith JN0919 (Fig. 7). However, due to their higher εHf(t), Groups II and III zircons are considered to have resulted from mixing between juvenile melts with older, late Archean crust, thus reflecting Paleoproterozoic underplating (Moyen et al., 2017). Here, we reexamine these two groups of zircons and compare them with Hannuoba counterparts. Group III zircons come from a mafic garnet granulite xenolith 02-34 and have the highest εHf(t) (–3.5 to –1.7). Their extremely low

176Lu/177Hf

values (0.000010–0.000031) are consistent with a

metamorphic origin involving garnet which is indeed present in the xenolith. They have εHf(t) values plotting roughly along the 2.7 Ga mafic crust evolution line with 176Lu/177Hf

= 0.022 (Fig. 5 of Moyen et al., 2017), similar to the ~1.8 Ga zircons for

Hannuoba mafic granulite JN0707 and amphibolite JN0731 which have εHf(t) values plotting along the 2.5 Ga mafic crust evolution line (Fig. 7). Importantly, Group III zircons have εHf(t) values (–3.5 to –1.7) lower than those (-1.0 to 2.9) of the ~1.8 Ga zircons for Hannuoba mafic granulite JN0707 and amphibolite JN0731. Therefore, these 1.8–1.9 Ga zircon ages most likely reflect metamorphism rather than

Paleoproterozoic underplating and their decreased THfDM ages of 2.3–2.4 Ga would be significantly younger than the protolith age for the xenolith. Indeed, many granulite xenoliths from Udachnaya contain ~2.7 Ga zircon ages, supporting a ~2.7 Ga protolith age for xenolith 02-34. As with Group II zircons, they plot slightly above the 2.7 Ga average crust evolution line with

176Lu/177Hf

= 0.015 (Fig. 5 of Moyen et al.,

2017), similar to the Mesozoic zircons from the Hannuoba mafic xenoliths (Fig. 7). They unlikely recrystallized from the pre-existing Archean grains as Group I zircons did. Instead, they are probably newly grown zircons either from metamorphic fluids or from melts derived from the pre-existing crust. In either case, they have a derivation from the Archean lower crust. To summarize, the three groups of 1.8–1.9 Ga zircons from the Udachnaya granulite xenoliths have origins perfectly comparable with zircons from the Hannuoba granulite terrain and xenoliths although their ages are different. Therefore, we consider that none of the reported 1.8–1.9 Ga zircons from Udachnaya reflect Paleoproterozoic underplating. If this is true, it does not support the suggestion by Moyen et al. (2017) that a complete or large-scale delamination and rejuvenation of the Archean lower lithosphere took place in the Paleoproterozoic. Recently, based on an integrated study of zircon U–Pb age and Hf isotopic compositions and whole rock Nd isotopic data, Shatsky et al. (2019) suggest that the 1.8–1.9 Ga zircons from lower crustal xenoliths of the Anabar Shield in the Siberian craton mainly reflect a tectono-thermal event and the upper and lower crust is coupled, rather than decoupled. When dealing with granulite xenoliths that were considered to be typical examples

resulting from basaltic underplating, the least and most isotopically evolved xenoliths are usually assumed to represent the mantle and crustal end-members, respectively, and the xenoliths with isotopic compositions in between are considered to result from mixing between mantle-derived basaltic melts and the pre-existing lower crust (Downes et al., 1990; Rudnick, 1990). However, as shown above, there is no evidence for the involvement of components from the most isotopically evolved xenoliths in the studied Hannuoba mafic xenoliths. Here we further suggest that the isotopically least evolved Hannuoba mafic xenoliths may not represent the mantle end-member either. On Fig. 4a, some samples from the granulite terrain in the Hannuoba region have positive εNd values (up to 17) due to time-integrated radiogenic growth given their high

147Sm/144Nd

(>0.19). Since the pre-existing lower crust beneath Hannuoba

could be represented by the granulite terrain, it is inferred that rocks with similarly depleted isotopic compositions may exist in the lower crust. Indeed, two Hannuoba garnet-bearing mafic xenoliths 90DA11 and 95SQ9 have εNd of 10.1 and 10.9, respectively (Fig. 4b) (Zhang et al., 1999). Precambrian ages and a 384 Ma zircon age (no Mesozoic ages) were reported for xenolith 90DA11 (Wilde et al., 2003). It is likely that this xenolith is a remnant of the pre-existing late Archean lower crust, rather than a product of young underplating. Therefore, xenoliths with depleted isotopic compositions are not necessarily products of young underplating. Our findings demonstrate that young zircon ages and a large range and/or depletion in Sr–Nd isotopic compositions of lower crustal xenoliths are insufficient criteria to infer recent basaltic underplating. It is implied that some granulite xenoliths

previously regarded to be products of basaltic underplating or crust–mantle mixing in many other regions based on zircon ages and/or Sr–Nd isotopic compositions may actually be remnants or derivatives of the pre-existing lower crust. If so, the role played by basaltic underplating in the evolution of the lower crust in these regions may need to be re-evaluated. Clearly, an integrated study of granulite terrain and xenoliths as well as surface magmatism can better constrain the evolution of the lower crust. Although most Hannuoba mafic xenoliths represent reworked, not the original, late Archean lower crust, the two xenoliths JN0808 and 95DA6 provide strong evidence for existence of original late Archean mafic lower crust. Mafic granulite xenoliths believed to be Archean in ages can also be found from the Kaapvaal craton (Huang et al., 1995), the Superior Province (Moser and Heaman, 1997), Man Shield of the Western African craton (Barth et al., 2002), the central and southern Slave craton (Davis et al., 2003), the Karelian craton (Peltonen et al., 2006), the northern Wyoming craton (Bolhar et al., 2007), the Tanzanian craton (Bellucci et al., 2011), the Churchill Province of Canada (Petts et al., 2014) and the Siberian craton (Shatsky et al., 2016, 2019). These findings demonstrate existence of widespread Archean mafic lower crust globally and argue against the generality by Rudnick and Gao (2014) that mafic granulite xenoliths from the lower crust of most Archean cratons have formed from post-Archean basaltic underplating. It is implied that processes in the formation of Archean crust may not be significantly different from those in present crust formation.

6. CONCLUSIONS The main conclusions of this study may be summarized as follows: (1) The Hannuoba Al-websterite xenoliths are most likely components of the lower crust, rather than samples from the mantle. They are cognate with the pyroxene-rich granulite xenoliths. (2) Both the Hannuoba Al-websterite and pyroxene-rich granulite xenoliths can be best explained as restites left after partial melting of the late Archean lower crust that may be represented by the granulite terrain to produce the voluminous Mesozoic intermediate–felsic magmatic rocks in the region. (3) Although a few 1.8–1.9 Ga and ~130 Ma zircons from the Hannuoba lower crustal xenoliths may record two periods of underplating, they are exotic to the xenoliths and the xenoliths are not products of either Paleoproterozoic or Mesozoic underplating. (4) Our findings demonstrate that high-grade metamorphism or partial melting of Archean rocks can result in zircons with significantly younger U–Pb (1.8–1.9 Ga or Mesozoic) and THfDM ages. The large range and/or depletion in Sr–Nd isotopic compositions of the lower crustal xenoliths could be inherited from the pre-existing lower crust. Therefore, young zircon ages and a large range and/or depletion in Sr–Nd isotopic compositions of lower crustal xenoliths are insufficient criteria to infer recent basaltic underplating. (5) Combined our results with literature data, it is concluded that Archean mafic lower crust is common globally, arguing against the previous suggestion that mafic granulite xenoliths from the lower crust of most Archean cratons might have formed from post-Archean basaltic underplating.

ACKNOWLEDGEMENTS Mary Horan and Tim Mock of DTM and Qian Mao, Chaofeng Li, Qiuli Li, and Yueheng Yang of IGGCAS are thanked for analytic and laboratory assistance. Constructive comments from R.W. Carlson on a preliminary draft of this paper are greatly appreciated. Three anonymous reviewers are thanked for detailed comments that helped us to improve the manuscript substantially. Dr. Y. Amelin is thanked for the editorial work. The research was financially supported by funds from the Ministry of Science and Technology of People's Republic of China (2016YFC0600103) and the National Natural Science Foundation of China (No.41973034).

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Table 1 Zircon U-Pb ages and Hf-O isotope compositions for Al-websterite xenolith HJ1435

Figure captions Fig. 1 Simplified geological map of the Hannuoba region showing the distribution of the southern outcrop of the xenolith-bearing Hannuoba basalts, the adjacent granulite terrain and the Mesozoic granitoids and volcanic rocks. Sample numbers and locations from the granulite terrain are shown. Also sample locations and their zircon ages for the Mesozoic intermediate–felsic igneous rocks are shown.

Fig. 2 Petrological and petrographic characteristics of Hannuoba pyroxene-rich granulite xenoliths and Al-websterite xenoliths. (a) Photomicrograph of pyroxene-rich granulite xenolith JN0808. (b) Hand specimen of dark black Al-websterite xenolith. (c) Photomicrograph of Al-websterite xenolith JN0901. (d) Photomicrograph of olivine-bearing Al-websterite xenolith JN0901. Cpx = clinopyroxene; Opx = orthopyroxene; Pl = plagioclase; Ol = olivine.

Fig. 3 REE patterns of various rocks from the Hannuoba region. (a) Rocks from the granulite terrain; (b) The pyroxene-rich granulite xenoliths; and (c) The Al-websterite xenoliths. Chondrite values are from Sun and McDonough (1989). Data are compiled in Tables S1–3.

Fig. 4 143Nd/144Nd vs. 147Sm/144Nd (a) and εNd–87Sr/86Sr (b) for various rocks from the Hannuoba region. The Sr–Nd isotope compositions of the 6 dated mafic xenoliths are shown in the upper right corner of Fig. 4b. Zircon ages: Al-websterite xenolith DMP0301B – exclusively Mesozoic with a mean of 158 ± 1 Ma (n = 5) (Wei et al., 2015); Al-websterite xenolith DMP-315 – dominantly Mesozoic (97–158 Ma) with additional ages of ~3.12 Ga, ~2.54 Ga, ~1.84 Ga, ~418 Ma and ~230 Ma (Liu et al., 2004); Al-websterite xenolith JN0737 – Archean with a mean of 2715 ± 21 Ma (n = 7) (Jiang et al., 2010); Al-websterite HJ1435 – 78–137 Ma (n = 10) with an additional age of 2136 Ma (This study,

Table 1); Pyroxene-rich granulite xenoliths JN0724 and JN0813 – ~2.5 Ga (Jiang and Guo, 2010; Jiang et al., 2011). The Sr–Nd isotopic compositions are compiled in Tables 1, S1–3, S5–7.

Fig. 5 Zircon CL image and data of Hannuoba xenoliths (a) Al-websterite xenolith HJ1435. Red circles in zircons show the position of SIMS U–Pb and O isotope analytical sites. The ages of each spot are 206Pb/238U ages (for Mesozoic ages) or 207Pb/206Pb

age (for Precambrian age of spot 12). The numbers below the ages

represent the corresponding O isotope compositions. The zircon CL images and O isotope compositions clearly show two groups of Mesozoic zircons. Group II zircons with O isotope compositions between 5.22 and 5.48‰ are relatively smaller in size and darker in CL images than Group I ones, which have higher O isotopic compositions (6.32–7.11‰) and are larger in size and brighter in CL images. (b) Intermediate granulite xenolith JN0919 from Jiang et al. (2011).

Fig. 6 Zircon U-Pb age histogram for granulite terrain, Al-websterite xenoliths, pyroxene-rich granulite xenoliths and Mesozoic magmatic rocks from the Hannuoba region. 207Pb/206Pb ages are employed for the Precambrian zircons and 206Pb/238U

ages are used for the Phanerozoic zircons. Data for granulite terrain are

from Jiang et al. (2010). Data for pyroxene-rich granulite xenoliths are from Wilde et al. (2003); Liu et al. (2004); Fan et al. (1998) and Jiang and Guo (2010). Data for Al-websterite xenoliths are from this study; Wilde et al. (2003); Liu et al. (2004); Jiang et al. (2010); Wei et al. (2015). Data for Mesozoic magmatic rocks

are from Jiang et al. (2011). All data plotted with discordance <5%, except that data from Wilde et al. (2003) where 207Pb/235U ages are not available.

Fig. 7 Whole rock and zircon Hf isotopic compositions for various rocks from the Hannuoba region. The

176Lu/177Hf

of 0.015 of average crust and of 0.022 of

mafic crust are from Griffin et al. (2002) and Amelin et al. (2000), respectively. To compare with the Mesozoic magmatic rocks, the whole rock εHf values of both the pyroxene-rich granulite and Al-websterite xenoliths are recalculated at 130 Ma. For clarity, they are plotted separately, not exactly at the position of t = 130 Ma. Zircon εHf(t) values are recalculated at their ages. Sample locations and zircon ages of the Mesozoic magmatic rocks are given in Fig. 1. Data are compiled in Tables 1, S2–3, and S8. Data for the pyroxene-rich granulite xenoliths are from this study; Liu et al. (2004); Jiang and Guo (2010); Jiang et al. (2011) and Hu et al. (2019). Data for Al-websterite xenoliths are from this study; Choi et al. (2008); Jiang et al. (2010); Hu et al. (2019) and Wei et al. (2019). Data for the granulite terrain are from Jiang et al. (2010). Data for 4 felsic granulite xenoliths are from Wei et al. (2015). Data for JN0919 are from Jiang et al. (2010). Data for Mesozoic granitoids are from Jiang et al. (2007, 2009, 2011).

Fig. 8 Lower crust normalized multi-element patterns for Mesozoic magmatic rocks, the granulite terrain, pyroxene-rich granulite xenoliths, and Al-websterite xenoliths from the Hannuoba region. Normalizing values are from Rudnick and

Gao (2014). Also listed are Sr–Nd isotopic compositions of the various rocks. Their (87Sr/86Sr)i and εNd(t) were recalculated at t = 130 Ma. Except for the granulite xenolith 95DA6 and trachyandesite sample JN0713, the remaining rocks are averages of currently available data. The data of the granulite terrain exclude samples with SiO2>60 wt.%. Data are compiled in Tables S1–3 and S6.

Fig. 9 (La/Yb)N and TNdDM vs. 147Sm/144Nd for Hannuoba pyroxene-rich granulite and Al-websterite xenoliths. Data are compiled in Tables S2–3.