Formation and alteration of a zoned calcsilicate vein from the contact aureole of Kasuga area, central Japan: insights from thorianite and uraninite chemical ages

Abstract

To understand fluid activity around a granitic pluton, a petrological study combined with U-Th-total Pb chemical dating of accessory U-Th oxides was conducted on a zoned calcsilicate vein from the contact aureole of the Cretaceous Kaizukiyama granite, Japan. The vein exhibits distinct zoning with the following sequence of major assemblages from the host rock (dolomitic marble) to the vein center: Zone 1 (comprising three mineral assemblages in order away from the host rock: clinohumite + calcite, forsterite + calcite, and diopside + forsterite + calcite), Zone 2 (clintonite + spinel + Al-rich clinopyroxene or pargasite + calcite), and Zone 3 (grossular + anorthite + clinopyroxene + calcite). Additionally, various U-, Th-, and REE-rich accessory minerals occur in Zones 2 and 3. Grossular and anorthite in Zone 3 are extensively replaced by later-stage minerals (clinozoisite or prehnite + muscovite). Zone 1 is likely a metasomatic reaction vein formed by interactions between the host dolomitic marble and aqueous silica in the infiltrating fluid. The mineral assemblage in Zone 2 likely resulted from interactions between Zone 1 and a fluid enriched in alumina. Thorianite grains in Zone 2 yield a chemical age of 97.0 ± 1.1 Ma, contemporaneous with the emplacement age of the Kaizukiyama granite. The primary minerals in Zone 3 probably formed during a crack-sealing process associated with new fluid infiltration, which also slightly interacted with Zone 2. The later-stage minerals in Zone 3 are attributed to low-temperature hydrothermal alteration. Uraninite grains in Zone 3 yield a chemical age of 88.1 ± 0.8 Ma, likely related to the hydrothermal alteration stage. Combining these results with the previously established cooling history of the Kaizukiyama granite suggests that low-temperature hydrothermal activity may have continued for a long period (∼ 6 Myr) in the surrounding area after the rapid initial cooling of the pluton.

INTRODUCTION

Carbonate country rocks around an igneous pluton serve as excellent recorders of fluid-rock interaction during contact metamorphism. In particular, dolomitic limestones are well-suited for evaluating fluid compositions and metasomatic processes due to their chemical simplicity and the significant variation in metasomatic mineral assemblages (e.g., Suzuki, 1977; Rice, 1979; Ferry, 1994; Bucher, 1998). Besides the major silicate minerals present in metasomatized dolomitic marbles (skarns and calcsilicate veins), accessory minerals provide valuable insights into metasomatic processes, particularly regarding the behavior of incompatible trace elements, including actinides (U and Th), high-field strength elements (HFSE), and rare earth elements (REE) (e.g., Gieré, 1986). U- and Th-rich accessory minerals are also useful for constraining the ages of fluid activity. Specifically, U-Th oxide minerals (uraninite and thorianite) are ideal targets for U-Th-total Pb chemical dating using an electron microprobe (e.g., Kotzer and Kyser, 1993; Cocherie and Legendre, 2007). The high U and Th contents, coupled with minimal initial Pb in these minerals, may enable high-precision chemical dating, even for relatively young geological events.

This study focuses on a zoned calcsilicate vein in dolomitic marble from the contact aureole of the Kaizukiyama granite pluton in central Japan. The studied vein is unique in three aspects: the local concentration of alumina, indicated by clintonite-bearing paragenesis; the presence of various U-, Th-, and REE-rich accessory minerals; and multiple generations of minerals in the central part of the vein. Clintonite [Ca(Mg_1+xAl_2−x)(Si_xAl_4−x)O₁₀(OH)₂, x varies around unity] is a Si-poor/Al-rich member of the mica group. Since the thermodynamic stability field of clintonite is relatively large (Rice, 1979), its rarity is the consequence of a specific chemical environment of Ca, Mg, Al-rich, and K-poor conditions. The primary objectives of this work are twofold: (1) to infer the formation process of the zoned calcsilicate vein, particularly the enrichment of alumina during fluid-rock interaction and the extent of later-stage overprinting of vein mineralogy, and (2) to constrain the age and duration of fluid activity around the granite pluton through the application of U-Th-total Pb chemical dating of thorianite and uraninite in the vein. Abbreviations for minerals and end-member components follow Warr (2021).

GEOLOGICAL SETTING

The Kaizukiyama granite represents a Cretaceous granitic pluton in central Japan (Fig. 1). The pluton is composed mainly of massive hornblende-bearing biotite granite and granodiorite (Saito and Sawada, 2000). The Kaizukiyama granite shows a Rb-Sr whole rock isochron age of 96.4 ± 4.8 Ma (Sawada et al., 1994), which is indistinguishable from biotite K-Ar ages of 98.8 ± 4.9 and 94.6 ± 4.7 Ma (Saito and Sawada, 2000). Al-in-hornblende geobarometry on the Kaizukiyama granite suggests an emplacement pressure of around 200 MPa (Hong Mei et al., 2019).

Figure 1. Location of the Kaizukiyama granitic pluton and Kasuga area in central Japan (Saito and Sawada, 2000; Enami et al., 2021). Isograds of the contact aureole in the Kasuga area are based on mineral assemblages in dolomitic marble (Suzuki, 1977). The sample locality is at the Kawai Pit.

The contact aureole around the Kaizukiyama granite has been studied in the Kasuga area, southern side of the granite (Fig. 1). In the Kasuga area, the contact aureole develops in the Funafuseyama Unit of the Jurassic Mino-Tamba Complex (e.g., Enami et al., 2021). The Funafuseyama Unit in this area contains substantial amounts of altered basalt, limestone, and dolomitic limestone as blocks of mudstone-matrix mélange (Suzuki, 1975; Saito and Sawada, 2000). Based on the mineral assemblages in dolomitic marble, the aureole has been divided into four zones: the talc zone, tremolite-calcite zone, diopside-dolomite zone, and forsterite-diopside zone (Fig. 1), in ascending order of metamorphic grade (Suzuki, 1977). The peak temperature estimates of the forsterite-diopside zone increase toward the contact with the pluton, ranging from 595 to 640 °C or more (Suzuki, 1977; Enami et al., 2021).

SAMPLE DESCRIPTION

Mineralogical zoning of a calcsilicate vein

A zoned calcsilicate vein in dolomitic marble (Figs. 2a and 2b) was collected from the Kawai Pit of the Kasuga mine, located in the forsterite-diopside zone (Fig. 1). The host dolomitic marble is a massive rock composed of dolomite and calcite, with no quartz or silicate minerals. The calcsilicate vein is ∼ 4 cm wide and consists of three mineralogical zones (Zones 1-3, in order from the host rock side) (Fig. 2c). Both the host dolomitic marble and Zone 1 exhibit a fine-grained granoblastic texture (Fig. 3). A unique feature of the vein is the presence of large clintonite crystals and Al-rich minerals in Zone 2 (Figs. 2c and 2d). Calcite is present throughout the vein, whereas quartz is absent in all zones. Vugs filled by monocrystalline calcite are observed in the central part of Zone 3 (Figs. 2b and 2e). The following microstructural descriptions are based on optical microscopy and backscattered electron (BSE) imaging.

Figure 2. Zoned calcsilicate vein from the Kasuga area. [(a) and (b)] Scans of polished vein samples. (c) Illustration of the mineralogical zones within the vein. (d) Photomicrograph (crossed polarized light) showing radial clintonite crystals. (e) Euhedral anorthite in a calcite-filled central vug. Calcite was partially removed using dilute HCl. An, anorthite; Cal, calcite; Chu, clinohumite; Cln, clintonite; Cpx, clinopyroxene; Cz, clinozoisite; Di, diopside; Dol, dolomite; Fo, forsterite; Grs, grossular; Ms, muscovite.

Figure 3. Backscattered electron images showing an outer part of Zone 1 and host dolomitic marble. Subzone boundaries within Zone 1 are indicated by yellow dashed lines.

Zone 1

This zone is subdivided into three subzones based on mineral assemblage: clinohumite subzone (Chu + Cal), forsterite subzone (Fo + Cal), and diopside subzone (Fo + Di + Cal) in order from the host rock side (Fig. 3). The clinohumite and forsterite subzones are narrow (typically several mm in thickness) calcite-dominant zones with disseminated grains of clinohumite or forsterite. The diopside subzone is typically ∼ 1 cm thick. It is composed of colorless equant diopside crystals (>80 vol%) with disseminated forsterite grains (Fig. 2a) and interstitial calcite.

Zone 2

This narrow discontinuous zone develops along the inner side of Zone 1 and is characterized by large (up to 5 mm) platy crystals of clintonite. Along the boundary with Zone 1, diopside grains with similar characteristics as those in Di subzone (Zone 1) are crosscut by brighter Zone 2 clinopyroxene in BSE images (Fig. 4a). Clintonite crystals show radial aggregates that extend towards the vein center (Figs. 2d and 4b), and are closely associated with green spinel and colorless clinopyroxene or greenish-blue pargasite under plane-polarized light. Pargasite occurs as coarse anhedral crystals that enclose cuspate clinopyroxene grains (Fig. 4a). Both clinopyroxene and pargasite are present around clintonite and in direct contact with it. However, only the former also occurs as inclusions within clintonite (Fig. 4b). Calcite fills interstitial spaces of these minerals. Along the boundary with Zone 3, clintonite crystals are always replaced by chlorite (Fig. 4b). Accessory minerals in this zone include thorianite (Figs. 4c and 4d), zirconolite (Fig. 4d), baddeleyite (Fig. 4e), REE-bearing apatite (Fig. 4f), and REE-rich epidote-group minerals (Fig. 4g). These accessory minerals occur as small (<50 µm) euhedral grains.

Figure 4. Backscattered electron images of Zones 2 and 3. (a) Compositional heterogeneity of clinopyroxene grains in outermost part of Zone 2. Inset shows cuspate clinopyroxene within pargasite. (b) Radial clintonite crystals in Zone 2 in association with spinel and clinopyroxene. Clintonite is replaced by chlorite along the contact with Zone 3. (c) Thorianite inclusion in clintonite. (d) Clintonite aggregates associated with thorianite and zirconolite. (e) Baddeleyite crystals within clintonite. (f) Zoned apatite crystals within clintonite. (g) REE-rich epidote group minerals (allanite-dissakisite) in Zone 2. (h) Primary and later-stage minerals in Zone 3. Prehnite veinlets crosscut primary grossular. Anorthite is pseudomorphed into muscovite-prehnite-thomsonite aggregates except for inclusions in grossular. (i) Clinopyroxene in Zone 2 and outer part of Zone 3. (j) Clinopyroxene with a zircon crystal in inner part of Zone 3. (k) Later-stage garnet, prehnite, and epidote-group minerals crosscut primary grossular. Inset right shows thorianite inclusion in grossular. (l) Clinozoisite/prehnite-muscovite aggregates replacing grossular in Zone 3. (m) Uraninite grains associated with REE-rich epidote in Zone 3. A beam damage after microprobe analysis is visible on the uraninite grain (inset left).

Zone 3

This zone occupies the central part of the vein, and its thickness substantially varies due to the irregular inner surface of Zone 2 (Fig. 2c). Based on crosscutting relationships and pseudomorphic textures, primary and later-stage minerals can be identified in Zone 3. The primary minerals include grossular, anorthite, clinopyroxene, and calcite (Figs. 2a and 2b). Primary grossular is almost colorless under plane-polarized light and occurs as coarse anhedral crystals (Figs. 2a and 2b). Anorthite occurs as inclusions in grossular (Fig. 4h), fine-grained aggregates (Fig. 2b), or coarse euhedral crystals in calcite-filled vugs along the center of Zone 3 (Figs. 2b and 2e). Clinopyroxene in this zone occurs as light green prismatic crystals. In BSE images of the boundary with Zone 2, darker clinopyroxene grains are partially overgrown by brighter clinopyroxene (Fig. 4i).

Later-stage minerals are fine-grained and occur as veinlets crosscutting colorless grossular or pseudomorphs after anorthite. The primary colorless grossular is crosscut by reddish garnet (brighter in BSE images), prehnite, and clinopyroxene in inner parts of Zone 3 (Figs. 4j and 4k). Along the boundary with Zone 2, colorless grossular is partially replaced by fine aggregates of acicular clinozoisite (identified by Raman spectroscopy) and muscovite (Figs. 4b and 4l). This clinozoisite-bearing domain appears pinkish in hand-specimen (Fig. 2a) and is in direct contact with chloritized rims of clintonite in Zone 2 (Fig. 4b). Primary anorthite crystals in Zone 3 are partially or completely replaced by muscovite + prehnite ± zeolite (thomsonite-Ca) aggregates (Fig. 4h). The secondary Ca-Al silicates in Zone 3 show a distinct distribution with prehnite in inner parts and clinozoisite in outer parts (boundary with Zone 2), but there is no crosscutting relationship between these minerals.

Accessory minerals in this zone include REE-rich epidote-group minerals, apatite, titanite, zircon, thorianite, and uraninite. Apatite and thorianite occur as inclusions in primary grossular (Fig. 4k). REE-bearing epidote-group minerals occur as disseminated grains within clinozoisite aggregates or prismatic crystals within veinlets crosscutting grossular (Fig. 4k). Uraninite occurs as small (<30 µm) subhedral grains associated with REE-rich epidote (Fig. 4m). Zircon occurs as brown short prismatic crystals (∼ 100 µm long) in inner parts of Zone 3. The zircon crystals display fine oscillatory zoning and locally show porous texture with abundant uraninite particles (Fig. 4j).

ANALYTICAL METHODS

Mineral chemistry

Chemical compositions of minerals were measured using a JEOL JXA-8530F electron microprobe at Shimane University. Operating conditions were 15 kV accelerating voltage (20 kV for REE-bearing minerals), 12-20 nA specimen current and a beam diameter of 1-4 µm. Well-characterized materials including oxides (Si, Ti, Al, Fe, Mn, and Zr), silicates (Mg, Ca, Na, and K), sphalerite (Zn and S), Y-Al garnet (Y), fluorite (F), apatite (P), boracite (Cl), metals (Hf, Nb, and Ta), and synthesized phosphates (REE) were used as standards. The wavelength-dispersive spectra were obtained for representative samples of REE-bearing minerals, and the peak (Lα lines for La, Ce, Dy, Er, Yb, and Lβ lines for Nd, Pr, Sm, Gd) and background positions were carefully determined. Representative analyses are reported in Tables 1 and 2.

Table 1. Representative analyses of major constituent minerals in the Kasuga calcsilicate vein

Mineral	Chu	Fo	Di	Cpx	Cln	Spl	Prg	An	Grs	Grt	Cpx	Cz	Prh
Zone	1	1	1	2	2	2	2	3	3	3	3	3	3
SiO2 wt%	38.71	40.62	54.64	50.19	16.79	n.d.	40.76	43.18	39.92	38.62	50.74	40.05	44.07
TiO2	n.d.	n.d.	n.d.	0.25	0.04	n.d.	0.28	n.d.	0.10	0.12	0.03	n.d.	0.03
Al2O3	n.d.	n.d.	1.21	6.95	43.25	64.03	18.50	35.47	21.08	19.97	0.31	33.27	24.37
FeOt	3.90	9.70	1.28	3.82	3.31	18.70	5.69	n.d.	2.80	12.85	19.10	0.08	0.03
MnO	0.33	1.05	0.13	0.18	0.05	0.89	0.16	n.d.	0.61	6.31	0.89	0.01	0.02
ZnO	-	-	-	-	-	1.84	-	-	-	-	-	-	-
MgO	55.45	49.54	17.19	13.42	18.39	15.14	14.92	n.d.	0.82	0.18	5.34	0.01	n.d.
CaO	0.02	n.d.	25.69	24.50	12.40	n.d.	13.08	20.31	33.99	22.27	23.66	24.58	27.09
Na2O	n.d.	n.d.	n.d.	n.d.	0.15	-	2.85	0.03	n.d.	n.d.	0.08	n.d.	n.d.
K2O	n.d.	n.d.	n.d.	n.d.	n.d.	-	0.57	n.d.	n.d.	n.d.	n.d.	n.d.	n.d.
F	2.99	-	-	-	0.27	-	0.46	-	-	-	-	-	-
Cl	0.01	-	-	-	n.d.	-	0.07	-	-	-	-	-	-
-O=F+Cl	1.26	-	-	-	0.11	-	0.21	-	-	-	-	-	-
Total	100.15	100.91	100.14	99.31	94.55	100.60	97.13	98.99	99.32	100.32	100.19	98.00	95.61
O=	-	4	6	6	11	4	23	8	12	12	6	12.5	11
Si	4.03	0.99	1.98	1.85	1.20	0.00	5.90	2.02	3.03	3.01	2.00	3.03	3.02
Ti	0.00	0.00	0.00	0.00	0.00	0.00	0.03	0.00	0.00	0.01	0.00	0.00	0.00
Al	0.00	0.00	0.05	0.30	3.65	1.96	3.16	1.96	1.89	1.84	0.01	2.97	1.97
Fe3+	0.00	0.00	0.00	0.00	0.00	0.04	0.00	0.00	0.04	0.12	0.00	0.01	0.00
Fe2+	0.34	0.20	0.04	0.12	0.20	0.36	0.69	0.00	0.14	0.72	0.63	0.00	0.00
Mn	0.00	0.02	0.00	0.01	0.00	0.02	0.02	0.00	0.04	0.42	0.03	0.00	0.00
Zn	-	-	-	-	-	0.04	-	-	-	-	-	-	-
Mg	8.60	1.80	0.93	0.74	1.96	0.58	3.22	0.00	0.09	0.02	0.31	0.00	0.00
Ca	0.00	0.00	1.00	0.97	0.95	0.00	2.03	1.02	2.77	1.86	1.00	1.99	1.99
Na	0.00	0.00	0.00	0.00	0.02	-	0.80	0.00	0.00	0.00	0.01	0.00	0.00
K	0.00	0.00	0.00	0.00	0.00	-	0.11	0.00	0.00	0.00	0.00	0.00	0.00
F	0.98	-	-	-	0.06	-	0.21	-	-	-	-	-	-
Cl	0.00	-	-	-	0.00	-	0.02	-	-	-	-	-	-
Cations	13.00	3.01	4.00	3.99	7.99	3.00	15.95	5.00	8.00	8.00	3.99	7.99	6.98

An, anorthite; Chu, clinohumite; Cln, clintonite; Cpx, clinopyroxene; Cz, clinozoisite; Di, diopside; Fo, forsterite; Grs, grossular; Grt, garnet; Prg, pargasite; Prh, prehnite; Spl, spinel. n.d., not detected.

Table 2. Representative analyses of accessory minerals in the Kasuga calcsilicate vein

Mineral	Ap	Ap	Ap	Aln	Aln	Aln	Aln	Bdy	Bdy	Zrc	Zrc	Zrn
Note	Core	Mantle	Rim		Mg-rich
Zone	2	2	2	2	2	3	3	2	2	2	2	3
SiO2 wt%	1.45	4.23	0.37	33.49	32.09	31.30	31.76	n.d.	n.d.	0.02	0.19	31.12
TiO2	-	-	-	n.d.	0.09	0.22	0.17	0.04	n.d.	30.58	29.24	n.d.
Al2O3	-	-	-	21.49	20.49	18.94	19.53	0.03	0.04	0.65	0.72	n.d.
FeOt	0.06	0.05	0.12	5.60	5.41	11.15	8.84	0.16	0.19	5.85	6.15	0.11
MnO	0.05	0.03	0.03	0.15	0.09	0.31	0.11	n.d.	n.d.	0.46	0.64	n.d.
MgO	0.01	0.00	0.04	2.72	3.15	0.51	1.55	0.03	0.03	0.12	0.10	-
CaO	52.70	48.35	54.76	14.88	13.46	13.99	14.32	0.05	0.01	10.82	9.39	0.01
Na2O	0.00	0.00	0.00	-	-	-	-	-	-	-	-	-
ZrO2	-	-	-	-	-	-	-	95.36	96.61	30.52	30.25	61.76
HfO2	-	-	-	-	-	-	-	3.61	3.30	1.32	1.25	2.63
Nb2O5	-	-	-	-	-	-	-	0.67	0.48	6.25	6.64	n.d.
Ta2O5	-	-	-	-	-	-	-	-	-	0.80	1.04	-
Y2O3	1.40	3.08	0.19	0.00	0.13	0.27	0.27	n.d.	n.d.	7.18	7.89	0.96
La2O3	0.12	0.17	n.d.	5.90	4.95	4.28	4.02	-	-	-	-	-
Ce2O3	0.35	0.57	0.09	7.75	8.03	7.13	6.62	0.15	0.02	0.10	0.14	0.05
Pr2O3	n.d.	0.09	n.d.	0.96	1.27	0.96	0.93	-	-	n.d.	0.12	-
Nd2O3	0.28	0.66	n.d.	2.22	3.68	2.19	2.18	n.d.	n.d.	0.22	0.24	n.d.
Sm2O3	0.09	0.21	n.d.	n.d.	0.25	n.d.	0.11	-	-	0.08	0.19	n.d.
Gd2O3	0.18	0.55	0.01	0.09	0.18	0.08	0.21	n.d.	n.d.	0.67	0.78	n.d.
Dy2O3	0.11	0.37	n.d.	n.d.	n.d.	0.06	n.d.	n.d.	n.d.	1.18	1.35	0.04
Er2O3	0.02	0.17	n.d.	-	-	-	-	n.d.	n.d.	0.85	0.90	0.12
Yb2O3	0.06	0.09	0.02	-	-	-	-	n.d.	n.d.	0.67	0.80	0.13
ThO2	0.08	2.95	0.07	1.23	1.59	4.00	4.55	-	-	0.17	0.19	0.57
UO2	-	-	-	-	-	-	-	0.01	0.08	-	-	2.50
P2O5	38.53	33.71	40.74	-	-	-	-	-	-	-	-	n.d.
SO3	0.01	0.01	0.01	-	-	-	-	-	-	-	-	-
F	3.42	3.74	3.74	-	-	-	-	-	-	-	-	-
Cl	0.08	0.33	0.07	-	-	-	-	-	-	-	-	-
-O=F+Cl	1.46	1.65	1.59	-	-	-	-	-	-	-	-	-
Total	97.54	97.71	98.69	96.48	94.86	95.39	95.17	100.11	100.76	98.51	98.21	100.00

Ap, apatite; Aln, allanite; Bdy, baddeleyite; Zrc, zirconolite; Zrn, zircon. n.d., not detected.

U-Th-total Pb chemical dating

Thorianite and uraninite grains were analyzed using the JXA-8530F microprobe. The operating conditions were 20 kV accelerating voltage, 50 nA specimen current and a beam diameter of 2 µm. The UMβ, ThMα, PbMα, YLα, CeLα, DyLα, FeKα, SiKα, PKα, and CaKα lines were measured on the samples. The PbMα line was counted 300 s on the peak and 150 s on the background positions with the PETH crystal. U-dominant samarskite (UMβ), thorianite (ThMα), and crocoite (PbMα) were used as primary standards. The interferences of ThMγ and ThM3-N4 on UMβ, and ThMζ_1,2 and YLγ on PbMα were corrected following to the procedure described in Suzuki and Kato (2008). The ZAF method was applied for matrix correction. Consideration of the counting errors in UO₂, ThO₂, and PbO analyses alone underestimates the error associated with single-spot chemical age. Following Cocherie and Legendre (2007), a minimum relative error of 2% (2σ) was applied to UO₂, ThO₂, and PbO determinations to account for other error sources such as standardization and matrix corrections. Results are reported in Supplementary Tables S1 and S2 (Supplementary Tables S1 and S2 are available online from https://doi.org/10.2465/jmps.240404). Uraninite grains in pegmatites from Naegi (Gifu Prefecture, Japan) and Kotoge (Fukuoka Prefecture, Japan), and U-rich thorite included in a monazite crystal in pegmatite from the Ishikawa district (Fukushima Prefecture, Japan) were simultaneously analyzed as secondary standards.

RESULTS

Major minerals in the calcsilicate vein and host dolomitic marble

Clinohumite and forsterite. Clinohumite grains in the Chu subzone (Zone 1) exhibit Mg# [= Mg/(Mg + Fe²⁺) atomic ratio] of 0.96-0.99 and F/(F + OH) of 0.49. Forsterite grains in the Fo and Di subzones (Zone 1) show Mg# of 0.87-0.93 and 0.87-0.91, respectively.

Clinopyroxene. The composition of diopside in the Di subzone (Zone 1) is characterized by low Al₂O₃ (0.5-2.4 wt%) and high Mg# (0.93-0.96) (Fig. 5a). Diopside compositions show no clear dependence on the distance to the subzone boundary. Clinopyroxene in Zone 2 contains significant amounts of Al₂O₃ (1.0-10.3 wt%) as Ca-Tschermak (Ca^[6]Al^[4]AlSiO₆) component. The ^[4]Al content in clinopyroxene increases with decreasing Mg# (0.77-0.96) (Fig. 5a). This compositional variation is associated with complex heterogeneity within grains (Fig. 4a). Clinopyroxene in Zone 3 shows a systematic compositional change with decreasing Al and Mg# from the outer (adjacent to Zone 2) to the inner parts of the zone (Fig. 5a). The clinopyroxene composition in the inner parts of Zone 3 is close to the binary diopside-hedenbergite solid-solution (Mg# = 0.33-0.68) (Fig. 5a). Individual clinopyroxene grains throughout Zone 3 also exhibit compositional heterogeneities, with lower Mg# for later generations (rim or crosscutting domains).

Figure 5. (a) Clinopyroxene and (b) garnet compositions in the calcsilicate vein.

Clintonite. Clintonite in Zone 2 shows compositional variation with Si = 1.19-1.24 apfu and Mg# = 0.89-0.94. Small amounts of fluorine in clintonite (F = 0.12-0.27 wt%) correspond to F/(F + OH) = 0.015-0.030.

Spinel. Spinel grains in Zone 2 are Al-rich and can be expressed as a ternary (Mg-Fe²⁺-Zn) solid solution: Sp_49-61Hc_36-43Ghn_3-14.

Ca-Al silicates. The composition of primary grossular in Zone 3 varies within Grs_83-90Alm_3-7Adr_1-5Sps_1-4Prp₃ (Fig. 5b). Later generations of garnet (brighter crosscutting domains in BSE images: Fig. 4k) are almost Mg-free and show a stepwise compositional change to lower grossular and higher almandine + spessartine contents: Grs₈₈Adr₆Alm₄Sps₂ to Grs₄₄Alm₃₀Sps₂₄Adr₂ (Fig. 5b). Anorthite grains in Zone 3 are chemically pure (An₁₀₀). Clinozoisite and prehnite in Zone 3 are almost Fe³⁺-free and very close to end-member compositions (Table 1).

Carbonates. In the host marble, calcite grains in association with dolomite (Mg# = 0.99) contain 8.5-10.3 mol% MgCO₃ (Mg# = 0.98-0.99). Calcite compositions within the calcsilicate vein (dolomite-free) show a systematic decrease in MgCO₃ content (mol%) and Mg# towards the vein center: Chu subzone (6.0-8.0 mol% and Mg# = 0.94-0.98), Fo subzone (2.2-3.6 mol% and Mg# = 0.87-0.94), Di subzone (1.7-2.6 mol% and Mg# = 0.83-0.94), Zone 2 (0.8-1.0 mol% and Mg# = 0.67-0.77), and Zone 3 (<0.6 mol% and Mg# < 0.43).

Others. Pargasite in Zone 2 shows a compositional range of Si = 5.89-6.30 apfu, ^[A](Na + K) = 0.77-0.92 apfu, Na/(Na + K) = 0.64-0.88, and Mg # = 0.77-0.85, and contains small amounts of fluorine (F = 0.28-0.54 wt%) (Table 1). Muscovite associated with clinozoisite or prehnite in Zone 3 shows a small compositional variation with Si = 3.04-3.18 apfu. The interlayer cations of muscovite are sorely occupied by potassium (K = 1.00 apfu).

U-, Th-, and REE-bearing accessory minerals in Zones 2 and 3

Apatite. In BSE images, euhedral apatite grains in Zone 2 show a concentric zoning with a bright annulus (Fig. 4f). The dark rim compositions are close to the ideal fluorapatite formula [Ca₅(PO₄)₃F], whereas the bright annulus exhibits high REE, Y, and Th contents (Table 2). The positive correlations between these elements and Si (Fig. 6) suggest the substitutions (REE³⁺,Y³⁺) + Si⁴⁺ = Ca²⁺ + P⁵⁺ and Th⁴⁺ + 2Si⁴⁺ = Ca²⁺ + 2P⁵⁺.

Figure 6. Plot of REE, Y, Th versus Si for apatite in Zone 2.

Epidote-group minerals. REE-rich epidote group minerals occur as chemically heterogeneous grains in Zones 2 and 3 (Figs. 4g and 4k). These grains in Zone 2 are higher in Mg/(Mg + Fe_total), and lower in Fe³⁺/Fe_total and ThO₂ than those in Zone 3 (Fig. 7 and Table 2). In Zone 2, REE-rich analyses correspond to Mg-rich allanite-(Ce) or Fe²⁺-rich dissakisite-(Ce) (Fig. 7). The Y content (proxy for heavy REE) in Mg-rich allanite-(Ce) in Zone 2 is lower than that in allanite-(Ce) in Zone 3 (Fig. 8).

Figure 7. Plots of REE + Y + Th versus Mg/(Mg + Fe_total) and Al for epidote-group minerals in Zones 2 and 3.

Figure 8. Chondrite-normalized REE patterns of zirconolite, apatite, and allanite-(Ce) in the calcsilicate vein. Normalization values are taken from McDonough and Sun (1995).

Zr-rich minerals. Baddeleyite grains in Zone 2 (Fig. 4e) show a small compositional variation (ZrO₂ = 95.3-96.6 wt%, HfO₂ = 3.30-3.73 wt%, and Nb₂O₅ = 0.37-0.67 wt%). REE contents in baddeleyite are mostly below the detection limits (Table 2). Zirconolite (ideal formula: CaZrTi₂O₇) in Zone 2 (Fig. 4d) contains appreciable amounts of Y and heavy REE (Fig. 8). The mean of zirconolite analyses (n = 10) gives an empirical formula:

  

\begin{align*} &\text{(Ca$_{0.69}$Y$_{0.26}$REE$_{0.09}$Mn$_{0.03}$Mg$_{0.02}$)$_{\Sigma 1.09}$}\\& \text{(Zr$_{0.93}$Hf$_{0.02}$)$_{\Sigma 0.95}$}\\& \text{(Ti$_{1.47}$Fe$_{0.32}$Nb$_{0.17}$Al$_{0.06}$Ta$_{0.02}$)$_{\Sigma 2.04}$O$_{7}$} \end{align*}

Oscillatory zoned zircon grains in Zone 3 (Fig. 4j) consist of U-rich brighter zones (up to 3.3 wt% UO₂) and U-poor darker zones or domains with uraninite particles.

Th-U minerals. In BSE images, thorianite and uraninite grains exhibit faint concentric zoning (Figs. 4d and 4m). Some uraninite grains display a darker Th-rich domain crosscut by brighter U-rich domains (Fig. 4m). Two thorianite grains in Zone 2 exhibit Th/(Th + U) atomic ratios of 0.86 (Fig. 4c) and 0.61-0.69 (Fig. 4d), respectively. Thorianite included in grossular in Zone 3 (Fig. 4k) shows high Th/(Th + U) ratios of 0.94-0.98. Four uraninite grains in Zone 3 show Th/(Th + U) = 0.04-0.18.

U-Th-total Pb chemical dating

The apparent ages of thorianite and uraninite were calculated by assuming an initial PbO = 0, because the small variations in UO₂* contents in those minerals (Fig. 9) render them unsuitable for the chemical isochron method (CHIME).

Figure 9. Plot of PbO versus UO₂* for thorianite and uraninite in the calcsilicate vein. UO₂* represents the corrected UO₂ content (UO₂ plus the equivalent of ThO₂). Two slopes show reference isochrones. Weighted mean ages are shown with rank-order plots of single-spot apparent ages (boxes with vertical lengths corresponding to 2σ uncertainties). IsoplotR (Vermeesch, 2018) was used to construct the plots.

Nine analyses of the two thorianite grains in Zone 2 (Figs. 4c and 4d) yielded single-spot apparent ages ranging from 100.1 ± 3.5 to 94.9 ± 3.3 Ma. The arithmetic mean and weighted mean ages of the thorianite analyses are 97.1 ± 3.7 (2σ) and 97.0 ± 1.1 (2σ) Ma (MSWD = 1.2), respectively (Fig. 9). On the other hand, four analyses of one thorianite grain in Zone 3 (Fig. 4k) gave younger apparent ages ranging from 94.6 ± 3.3 to 87.7 ± 3.0 Ma. 14 spot analyses of four uraninite grains in Zone 3 resulted in single-spot apparent ages ranging from 91.1 ± 3.2 to 85.8 ± 3.0 Ma. The arithmetic mean and weighted mean ages of the uraninite analyses are 88.1 ± 3.0 (2σ) and 88.1 ± 0.8 (2σ) Ma (MSWD = 0.96), respectively (Fig. 9). The weighted mean ages are adopted in later discussion.

Simultaneous analyses of the secondary standards provided weighted mean ages of 105.3 ± 1.6 Ma (n = 5) for U-rich thorite from the Ishikawa pegmatite, 72.1 ± 1.0 Ma (n = 6) for uraninite from the Naegi pegmatite, and 95.7 ± 1.4 Ma (n = 6) for uraninite from the Kotoge pegmatite. The apparent age of the U-rich thorite agrees with the CHIME age (103 ± 26 Ma) of monazite from the same locality (Banno, 2021) as well as the zircon U-Pb age (104.2 ± 0.7 Ma) of the Ishikawa pluton (Takahashi et al., 2016). The Naegi uraninite age is also consistent with the zircon U-Pb age (71.3 ± 1.6 Ma) of the Naegi granite (Nakajima et al., 1993). Similarly, the apparent age of the Kotoge uraninite aligns with the apparent age (96.8 Ma) calculated from wet chemical analysis of uraninite from the same locality (Kimura and Iimori, 1937). These results would assure the accuracy of U-Th-total Pb ages of the samples.

DISCUSSION

Petrogenesis of the zoned calcsilicate vein

In general, the formation mechanisms of symmetrically-zoned calcsilicate veins in marble can be classified into three different types: 1) metasomatic reaction veins formed by reactions between minerals in the wall rock and dissolved species in infiltrated fluid along a crack (e.g., Bucher, 1998), 2) crack-seal veins formed by precipitation of dissolved components in infiltrated fluid, and 3) planer reaction zones formed by reactions between a basic dike or layer and host carbonate rock (e.g., Fukuyama et al., 2006). The first two types require a large fluid flux along an open fracture or permeable zone, whereas the last type presumes a pre-existing rock. The studied vein shows evidence of a formerly open fracture, such as radial clintonite crystals grown towards the vein center in Zone 2 (Fig. 2d), and calcite-filled vugs in the central part of Zone 3 (Fig. 2e). In the following discussion, the formation of the studied vein is explained by a combination of metasomatic reactions (Zones 1 and 2) and a crack-sealing process (Zone 3).

Zone 1. The presence of calcite and the absence of quartz are common to both the host dolomitic marble and all zones of the calcsilicate vein. The chemical system of Zone 1 can be approximated as a pseudo-binary (Mg,Fe)O-SiO₂ system in the presence of excess calcite and H₂O-CO₂ fluid (Fig. 10). In this context, the host dolomitic marble corresponds to the (Mg,Fe)O end-member of the system. The main features of Zone 1 include the sharp boundaries (parallel to the vein-host rock interface) of the Chu, Fo, and Di subzones (Fig. 3), along with a progressive increase in the SiO₂/(Mg,Fe)O ratio of the subzone system composition in this order (away from the host dolomitic marble). The preservation of chemical gradients across Zone 1 is consistent with a metasomatic reaction vein formed by a diffusion-controlled process. The net reaction for Zone 1 formation can be written as:

  

\begin{equation} \text{Dol} + \text{SiO$_{2}$ aq} = (\text{Di} + \text{Fo} + \text{Cal})_{\text{Zone1}} + \text{CO}_{2} \end{equation} (1).

Silica is one of the major components in magmatic/hydrothermal fluids derived from a granitic pluton. Therefore, the governing process associated with Zone 1 involves the diffusion of aqueous silica from the fluid in the central fracture to the reaction front in the dolomitic marble (e.g., Bucher, 1998). Diopside-forsterite veins with similar mineralogical compositions as Zone 1 are common around the sample locality and show crosscutting relationships, implying Si metasomatism of dolomitic marble along former fluid conduits (Enami et al., 2021), rather than the metasomatic replacement of chert layers.

Figure 10. Mineral compositions of Zones 1 and 2 in the calcsilicate vein plotted on a SiO₂-Al₂O₃-(Mg,Fe)O plane projected from calcite and H₂O-CO₂ fluid. Chu, clinohumite; Cln, clintonite; CaTs, Ca-Tschermak component; Di, diopside; Dol, dolomite; Fo, forsterite; Spl, spinel.

Zone 2. The primary mineral assemblage of Zone 2 is Al-rich clinopyroxene + spinel + clintonite + calcite, which can be expressed in the pseudo-ternary (Mg,Fe)O-SiO₂-Al₂O₃ system in the presence of excess calcite and H₂O-CO₂ fluid (Fig. 10). Al-rich clinopyroxene in Zone 2 locally crosscuts Al-poor diopside grains that have similar compositions to diopside in Zone 1 (Fig. 4a). This textural relationship suggests that the space now occupied by Zone 2 was formerly the Di subzone (Zone 1), and that minerals in the Di subzone acted as reactants for the formation of Zone 2. Therefore, a probable reaction forming Zone 2 can be written as:

  

\begin{align} &(\text{Di} + \text{Fo} + \text{Cal})_{\text{Zone1}} + \text{Al$_{2}$O$_{3}$ aq} + \text{H$_{2}$O} \\&\quad= (\text{Cpx} + \text{Spl} + \text{Cln})_{\text{Zone2}} + \text{CO}_{2} \end{align} (2).

The discontinuous distribution of Zone 2 (Fig. 2c) may be a consequence of the spaced distribution of reactant forsterite in Zone 1 (Fig. 2a). Subsequent crystallization of pargasite in Zone 2 (Fig. 4a) probably occurred via a reaction involving Al-rich clinopyroxene and alkalis in the fluid.

The mineral assemblage of Zone 1 is the same as the peak assemblage of the forsterite-diopside zone of the aureole (Fig. 1). The petrogenetic grid of Rice (1979) shows that the Cln-Cpx-Spl assemblage in Zone 2 is also stable under high temperature and high X_CO2 conditions. Therefore, the formation of Zones 1 and 2 is attributed to isothermal fluid-rock interaction during the peak of contact metamorphism (∼ 600 °C). The progression of Si metasomatism during Zone 1 formation may have caused chromatographic element fractionation, resulting in the depletion of SiO₂ relative to other dissolved components such as Al₂O₃ and alkalis in the fluid. The faster grain boundary diffusion of Si compared to Al could explain why Si metasomatism occurred first. After the formation of Zone 1, the reacted fluid likely had an elevated Al₂O₃/SiO₂ ratio in its dissolved components, and further migration of such fluids may have contributed to the formation of Zone 2 via reaction (2). The fluid associated with Zone 2 metasomatism was also enriched in incompatible trace elements, as evidenced by the occurrence of REE-rich apatite, allanite-dissakisite, thorianite, baddeleyite, and zirconolite. The presence of F-bearing silicates and apatite in Zone 2 suggests that fluorine in fluid may have enhanced the mobility of REE and HFSE by forming complexes (e.g., Gieré, 1986).

Zone 3. The primary process responsible for Zone 3 formation is the simultaneous crystallization of grossular and anorthite. Grossular in Zone 3 exhibits an overgrowth (without crosscutting) on the irregular surface of Zone 2, and contains no relics of Zone 2 minerals. The presence of euhedral anorthite within monocrystalline calcite (Figs. 2b and 2c) suggests that a substantial portion of the space now occupied by Zone 3 was formerly an open fracture, with minerals in Zone 3 essentially formed by precipitation from a fluid. However, clinopyroxene grains along the boundary between Zones 2 and 3 show an overgrowth texture of Fe-rich clinopyroxene on Al-rich clinopyroxene (Fig. 4i), making the exact position of the zone boundary difficult to define. This textural relationship can be interpreted either as the modification of Al-rich clinopyroxene in Zone 2 to Fe-richer compositions, or the formation of Fe-rich clinopyroxene in the outermost part of Zone 3 at the expense of Al-rich clinopyroxene in Zone 2. In either case, there is a minor but distinct interaction between Zone 2 and the fluid responsible for Zone 3 formation. Garnet-clinopyroxene-plagioclase veins with mineralogies similar to that of Zone 3 are pervasively developed within metasomatized basic rocks near the sample locality (Suzuki, 1975; Enami et al., 2021). If those veins are genetically linked to Zone 3, then Zone 3 was likely formed from a fluid that had previously passed through the basic rocks and been oversaturated in these mineral components. The stability field of the grossular-anorthite association is >470 °C at 200 MPa (Fig. 11). The lower temperature limit might decrease several tens of degrees when considering the reduced activity of grossular in the garnet composition within Zone 3. Thus, primary minerals in Zone 3 were formed during a period between the peak of contact metamorphism (∼ 600 °C) and an early stage of cooling (∼ 450 °C).

Figure 11. Isobaric (200 MPa) temperature versus X(CO₂) diagram showing the stability of the primary and later-stage assemblages in Zone 3 in the system CaO-Al₂O₃-SiO₂ (H₂O-CO₂ fluid in excess). Locations of two invariant points (IP1 and IP2) at 100 MPa are also shown. The petrogenetic grid was calculated using Perple_X_7.1.5 (Connolly, 2009) and the thermodynamic dataset (Holland and Powell, 1998). An, anorthite; Cal, calcite; Cz, clinozoisite; Dsp, diaspore; Grs, grossular; Prh, prehnite; Qz, quartz; Wrk, wairakite; Zo, zoisite.

Later-stage minerals in Zone 3 include chemically pure prehnite or clinozoisite. The prehnite-calcite association is stable between ∼ 280 and 420 °C at 200 MPa (Fig. 11). The lower temperature limit decreases to ∼ 250 °C at 100 MPa. In a Fe³⁺-free system, clinozoisite is a low-temperature polymorph of orthorhombic zoisite. The stability field of Fe³⁺-free clinozoisite is inferred to be sub-greenschist facies conditions from natural data (e.g., Enami and Banno, 1980) or below ∼ 370 °C from thermodynamic calculations (Fig. 11). Clinozoisite aggregates are distributed adjacent to Zone 2, although they never directly contact pristine Zone 2 minerals (Fig. 4b). There is no clear crosscutting relationship between prehnite and clinozoisite, and their spatial distribution in Zone 3 likely reflects local variation in silica activity:

  

\begin{equation} \text{Prh} + \text{CO}_{2} = \text{Cz} + \text{Cal} + \text{SiO$_{2}$ aq} + \text{H$_{2}$O}\end{equation} (3).

In summary, the formation of clinozoisite-muscovite aggregates (Fig. 4l), veinlets crosscutting primary grossular (Fig. 4k), and muscovite-prehnite pseudomorphs after anorthite in Zone 3 (Fig. 4h) as well as the chloritization of clintonite in Zone 2 (Fig. 4b) are related to retrograde hydrothermal alteration.

Interpretation of thorianite and uraninite ages

The rapid cooling of the Kaizukiyama granitic pluton has been inferred from the Rb-Sr whole rock and K-Ar biotite ages (Saito and Sawada, 2000). However, the present study reveals a significant difference in apparent ages between thorianite grains in Zone 2 and uraninite grains in Zone 3. Considering 2σ errors, the thorianite chemical age (97.0 ± 1.1 Ma) in Zone 2 overlaps with the Rb-Sr whole-rock age (96.4 ± 4.8 Ma), which probably represents the emplacement age of the pluton (Sawada et al., 1994). This supports the petrological inference that Zones 1 and 2 formed successively during the peak of contact metamorphism near the pluton. The oldest single-spot age of thorianite in Zone 3 (94.6 ± 3.3 Ma) may represent the crystallization age of primary Zone 3 minerals. In contrast, the uraninite chemical age (88.1 ± 0.8 Ma) and the rest of the thorianite single-spot ages in Zone 3 are significantly younger than the K-Ar biotite cooling ages in the Kaizukiyama granite (Saito and Sawada, 2000). Given the association of uraninite grains with later-stage epidote-group minerals in Zone 3 (Fig. 4m), the uraninite age may reflect crystallization during the retrograde hydrothermal stage or fluid-assisted resetting. While secondary Pb loss typically increases Ca and Si contents in uraninite (Kotzer and Kyser, 1993), such signs were not observed in the studied uranintie grains. U, Th, and Pb contents in the uraninite grains vary along the 88 Ma reference isochron (Fig. 9). Therefore, the uraninite age seems to be geologically meaningful and likely relates to the low-temperature (250-370 °C) hydrothermal stage in the aureole. Considering the closure temperature of the biotite K-Ar system at 300 ± 50 °C and the minimum bracket of the biotite K-Ar age at 94 Ma, low-temperature hydrothermal activity around the pluton may have continued ∼ 6 Myr after the rapid initial cooling of the granitic pluton. Similar plutonic cooling histories have shown initial rapid cooling (from the granitoid solidus to ∼ 300 °C) followed by slower cooling phases (several Myr/50 °C) associated with hydrothermal activity (e.g., Zhao et al., 2004; Yuguchi et al., 2011). Thus, the significant gap (∼ 9 Myr) between thorianite and uraninite ages does not contradict the proposed cooling history of the Kaizukiyama granite.

Beyond analytical uncertainties, unresolved geological issues such as non-zero initial PbO and isotopic discordance in thorianite and uraninite chemical ages underscore the need for further refinement of the cooling history of the Kaizukiyama granite and its aureole. Conducting a multi-chronological study, including in situ U-Pb dating of magmatic and hydrothermal zircon grains alongside other accessory minerals, is essential to elucidate the duration of fluid activity in the aureole.

ACKNOWLEDGMENTS

This study initiated when the author was a student at Nagoya University. At that time, he was impressed by the CHIME dating developed by the late Prof. K. Suzuki. I also thank emeritus Prof. A. Takasu for providing some standard materials. I am grateful to Drs. T. Ikeda and M. Uno for their constructive reviews of an earlier version of this manuscript, and to Prof. T. Kawakami for his editorial handling.

SUPPLEMENTARY MATERIALS

Supplementary Tables S1 and S2 are available online from https://doi.org/10.2465/jmps.240404.

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