DOI: 10.31038/GEMS.2025742

Abstract

Quaternary crustal movement of the Japanese islands caused the rapid uplift of mountain ranges and the relative subsidence of coastal and inland basins. In particular, from the Early Pleistocene, there was rapid uplift of the major mountain ranges, and in coastal areas, fan-deltas caused the burial of land shelf slopes and fan expansion. As one example, this paper presents the depositional processes of the Ogasa Group, which consists of gravelly fan-delta deposits on the Pacific coast in central Japan. The Ogasa Group divided into the Soga, Osuga, Kasui and Fukuroi Formations in ascending order. The depositional ages of these formations are estimated to be 1.78–1.19 Ma (Soga Formation), 1.19–0.91 Ma (Osuga Formation), 0.91–0.77 Ma (Kasui Formation), and after 0.77 Ma (Fukuroi Formation). The depositional processes and paleogeographical changes of the Ogasa Group can be divided into four stages as follows. Stage A: during the depositional stage of the Soga Formation, sand and mud were deposited from the shoreface to the continental slope, and at the late of this stage, the northwestward submarine channel was buried by coarse-grained sediments from the Tenryu River. Stage B: during the depositional stage of the Osuga Formation, the formation of a large fan-delta by the Oi River from the northeast, which successively advanced to the southwest, expanding the fan and extending offshore. Therefore, an inner bay was formed on its northwestern side, and sand and gravel from the Tenryu River were deposited. Stage C: during the depositional stage of the Kasui Formation, a lake was formed by sandbars formed offshore duo to the transgression. Stage D: during the subsequent deposition of the Fukuroi Formation, the fan formed by the Tenryu and Oi Rivers was thought to have expanded to the south.

Keywords

Middle Pleistocene, Ogasa Group, Fan-delta, Akaishi Mountains, Oi River, Tenryu River

Introduction

In central Honshu, Japan, there was rapid uplift of the Akaishi Mountains and other major mountain ranges from the late Pleistocene (Calabrian), and the land area expanded due to the supply of gravelly sediments in the surrounding areas [1]. Based on detailed stratigraphic and chronostratigraphic data of the Pliocene to Pleistocene in the central Honshu area, Shiba [2,3] divided the stratigraphic processes into four stages, placing the late early Pleistocene to early middle Pleistocene (Calabrian to early Chibanian), when the entire area became terrestrial with large-scale uplift of the Akaishi Mountains and other mountain ranges, in the Stage 2. During this stage, the island arc underwent a general uplift, and the subsidence area of the inland basin expanded, while in the coastal areas, the fan- delta caused the burial and landwardization of the shelf slopes and the expansion of fan areas. The Ogasa Group [4] is distributed in the coastal area of the Pacific Ocean between the Tenryu and Oi Rivers in western Shizuoka Prefecture (Figure 1) and consists mainly of gravelly fan-delta deposits from the late Early to early Middle Pleistocene. The large amount of sand and gravel comprising the Ogasa Group is supplied by the Tenryu and Oi Rivers, which originate from the Akaishi Mountains in the northern part of the distribution area, and the fan-deltas were formed when the Akaishi Mountains began to uplift on a large-scale [5]. Therefore, clarifying the details of the stratigraphy of the Ogasa Group and its depositional processes will help to better characterise the Stage 2 [2,3] of the Quaternary crastal tectonics. Shiba et al. [6] established the stratigraphy and depositional age of the Ogasa Group and inferred the depositional process based on the reconstruction of the depositional environment and its changes. This paper introduces the depositional processes of the Ogasa Group based on that paper [6].

Figure 1: Location map showing the study area indicated by the red box. R, River; Mts, Mountains; Mt. Mountain.

Stratigraphic Overview of the Ogasa Group

The Ogasa Group is Pleistocene series, mainly in Kakegawa and Fukuroi cities, and is exposed northwest to southeast on the Iwatahara Plateau, Enden Hills, Kasui Hills, Ogasa Hills and Minamiyama Hills (Figure 2). Many stratigraphic and palaeontological studies [4,7,8] have been carried out on the Ogasa Group. However, the lithology of the Ogasa Group is composed of thick gravels interbedded with sand and mud beds, which repeat from the lower to the upper part of the Group many times. The Ogasa Group is distributed far apart on the Ogasa Hills and other Hills, with different lithologies on each Hills, it was difficult to correlate the stratigraphy between hills.

Figure 2: Geological map of the Ogasa Group. M., Member.

The Ogasa Group consists of fan-delta deposits, mainly gravels, and unconformably overlies the uppermost part of the Kakegawa Group, the Hijikata Formation. The Ogasa Group is subdivided into the Soga, Osuga, Kasui and Fukuroi Formations (Figure 3), northwest-southeast running, inclined 5–10 degree to the southwest. The Kakegawa Group is a marine formation consisting of sand, mud, alternating sand and mud, and gravel beds, and is mostly composed of submarine fan deposits such as turbidite, except for shallow-water shelf deposits of the transgressive period in the upper part of this Group. The strata of the Kakegawa Group strike northwest-southeast and dip 15–30 degree to the southwest.

Figure 3: Stratigraphy and geological age of the Ogasa Group. F., Formation; M., Member; Strat., Stratigraphy. Cobb Mtn., Cobb Mountain. Wave line shows erosional surface.

The Soga Formation consists of the Aburayama sand Member in the Iwatahara Plateau, Enda Hills and Kasui Hills. In the Ogasa Hills, the Soga Formation consists of the Kogosho sand, Itasawa sand, Kamihijikata mud, Gansyoji gravel and Shinonba sand Members, while in the Minamiyama Hills, the Minamiyama gravel Member is distributed. The Osuga Formation is subdivided into the Takatenjin gravel, Ogasayama gravel, Oshiro sand, Hirano gravel, Obuchi gravel, Kamiishino sand, Higashioya sand and Mitsutoge gravel Membersin the Ogasa Hills. On the other hills, it consists of the Muramatsu gravel and mud Member. The Kasui Formation consists of the Kuno mud Member in the Kasui Hills and the Hattasan mud Member in the Ogasa Hills. The Fukuroi Formation consists of the Mukasa gravel Member in the Iwatahara Plateau, the Horikoshi gravel and mud Member in the Kasui Hills and the Yokosuka gravel Member on the Ogasa Hills.

Depositional Age of the Ogasa Group

The basement of the Ogasa Group is directly above the upper limit of the Olduvai Subchron [9], which corresponds to the Calabrian basement (approximately 1.8 Ma). Kameo [10] recognised the occurrence of the calcareous nannofossil Gephyrocapsa caribbeanica near the horizon of the Haruoka volcanic ash layer in the lower part of the Soga Formation at the eastern foot of the Ogasa Hills, and Gephyrocapsa oceanica slightly above the volcanic ash layer. According to Takayama et al. [11], the Haruoka volcanic ash layer was deposited at 1.72–1.65 Ma. Kameo [10] recognised the extinction of a large Gephyrocapsa genus in the uppermost horizon of the Soga Formation at the eastern foot of the Ogasa Hills, and assigned a base date of 1.2 Ma according to Takayama et al [11]. Ishida et al. [8] found normal magneticepoch in the uppermost part of the Soga Formation northwest of Takatenjin and in the basal part of the Gansyoji gravel Member. This is contrasted with the Cobb Mountain Subchron when combined with data from Kameo [10]. Therefore, the Soga Formation at the eastern foot of the Ogasa Hills is estimated to have been deposited between 1.78 and 1.19 Ma. In contrast, the Soga Formation of the Kasui Hills is not distributed strata in the Soga Formation above 1.65 Ma because its uppermost stratigraphic horizon is the Haruoka volcanic ash layer. The Soga Formation was deposited between 1.78 and 1.19 Ma, which means that the Osuga Formation was deposited after 1.19 Ma. Ishida et al. [8] recognised that the Kamiishino sand Member and the lower part of the Muramatsu gravel and mud Member of the Kasui Hills were deposited during the normal magnetic epoch, and compared their horizons to the Jaramillo Subchron. The Jaramillo Subchron is 1.07–0.99 Ma according to Gibbard and Head [12]. Therefore, in the Kasui Hills, the Muramatsu gravel and mud Member may have been deposited after 1.07 Ma, the base of the Jaramillo Subchron, while the upper part of this Member is reverse magnetic epoch, suggesting that it was deposited after 0.99 Ma. The upper part of the Osuga Formation in the Iwatahara Plateau is intercalated with the Kamikanzo volcanic ash layer. This ash layer is contrasted with the U8 volcanic ash layer of the Umegase Formation of the Kazusa Group, whose eruption age is estimated to be 0.92–0.91 Ma (MIS 23–24) according to Suzuki et al. [13].

The Kasui Formation of the Kasui Hills is intercalated with the Kunovolcanic ash layer, which is correlate to the Azuki volcanic ash layer, which is intercalated with the Ma 3 Member of the Osaka Group; Satoguchi and Nagahashi [14] estimated the age of this ash layer to be about 0.85 Ma. The Kasui Formation of the Ogasa and Kasui Hills was deposited in the reverse magnetic epoch [8], which corresponds to the horizon between the Jaramillo Subchron and the Brunhes Chron. The Kasui Formation is considered to have been deposited between 0.91–0.77 Ma based on the absence of Metasequoia and the inferred age of the lower Osuga Formation.

According to Ishida et al. [8], the four mud beds interbedded near the base of the Fukuroi Formation (the Yokosuka gravel Member) in the Ogasa Hills and the mud beds of the Horikoshi gravel and mud Member in the Kasui Hills are all normal magnetic epoch. This suggests that the base of the Fukuroi Formation was deposited after 0.77 Ma at the Brunhes-Matuyama Chron boundary, i.e. the base of Chibanian. The age of the uppermost part of the Ogasa Group cannot be specifically estimated because the uppermost horizon of the Ogasa Group could not be determined in this study. However, if the Ogasa Group is contrasted with the Osaka and Kazusa Groups, the uppermost part of the Ogasa Group is probably dated at around 0.4 Ma, which is considered to be the upper limit of both groups.

Depositional Environment of the Ogasa Group

The Soga Formation, the lowermost member of the Ogasa Group, is distributed in the Iwatahara-Kasui area by the Aburayama sand Member, which was deposited in the upper to lower shoreface. On the Ogasa Hills, on the other hand, the Kamihijikata mud Member was deposited on the continental slope from the outer continental shelf, while the Kogosho sand and the Itasawa sand Members were deposited on the inner shelf during the storm. The submarine channel deposits, the Gansyoji gravel and the Minamiyama gravel Members, were deposited above them, and the Shinonba sand Member, which have been deposited on the upper shoreface to inner shelf, accumulated. The sands of the Soga Formation include biotite and the gravels contain quartz porphyry and diorite, all of which are presumed to have been supplied by the Tenryu River system. In the Iwatahara-Kasui area, the Osuga Formation consists of the Muramatsu gravel and mud Formation, which is presumed to have been formed by several repetitions of the depositional environment of the inner bay or lagoon and upper shoreface or the diverging sandbar near the river mouth. From the sand and gravel of this member, it is assumed that these sediments were supplied by the Tenryu River system. On the other hand, the Osuga Formation at the eastern foot of the Ogasa Hills consists of the Takatenjin gravel Member, which was deposited near the mouth of the river from a sandbar on the river margin, followed by large-scale fan-delta front gravels (the Ogasayama gravel and Obuchi gravel Members) and fan braided river deposits (the Hirano gravel and Mitsutoge gravel Members). They consist of delta to fan deposits and are thought to have buried continental slopes. The absence of Tenryu River derived elements in the gravels of the Takatenjin gravel Member suggests that the sediments were supplied from the Oi River system and that the fan-delta of the Oi River system was first formed in the Ogasa Hills area during the Osuga Formation depositional period. The Ogasayama gravel Member is thought to have been supplied by the Oi River system, and this fan-delta expanded by advancing from north-east to south-west in the direction of the tilted stratification. The Oshiro sand Member overlying the Ogasayama gravel Member is thought to have been deposited on the upper shoreface, and the Hirano gravel Member of fan deposits and the Obuchi gravel Member deposited on the fan-delta front to the south of it were deposited above it. These gravel beds are also composed of sediments from the Oi River system, and this horizontal and vertical succession of lithologies corresponds to the progressive delta of Reading and Collinson [15], suggesting that the fan-delta developed in a more southerly offshore area at this time. The fan-delta is thought to have developed in a more southerly offshore extent at this period.

The Hiranogravel and the Obuchi gravel Members are overlain by the Kamiishino sand Member, the Higashioyasand and the Mitsutogegavel Members. These Members are deposits of lower shoreface, lower shoreface to backshore and fan, respectively. The presence of the Kamiishino sand and the Higashioya sand Members suggests that the transgression occurred prior to the deposition of these Members. The subsequent relative sea-level stagnation or falling of the sea-level formed the Kamiishino sand and Higashioyasand Members, and the fan sediments, the Mitsutoge gravel Member, was deposited above them, suggesting that an advancing fan-delta developed further to the south, extending the fan. The deposition of the Osuga Formation in the Ogasa Hills is thought to have completely buried and terrestrialised the continental slopes. In the sediments of the Osuga Formation, no sediments such as biotite sands or quartz porphyry gravels suggest that they were supplied from the Tenryu River system, except for the Kamiishino sand Member, and these sediments are thought to have been supplied from the Oi River system. The Kuno mud Member distributed in the Kasui Hills are presumed to have been deposited on the bottom of a freshwater lake with little influence from waves, based on the mud beds with well-developed laminae and shell fossils. The Hattasanmud Member in the Ogasa Hills is also presumed to have been deposited on a water bed based on its lithology. The Kasui Formation is considered to have been deposited during a transgressive period, because the depositional period of the Kasui Formation correlates with that of the Ma 3 Member, a transgressive deposit of the Osaka Group, and the mud beds of the Kasui Formation are widely distributed continuously from the Ogasa Hills to the Kasui Hills. Regarding the depositional environment of the Fukuroi Formation, the Mukasa gravel Member of the Iwatahara Plateau and the Yokosuka gravel Member of the Ogasa Hills are similar to the braided river depositional facies of Miall [16] and are considered braided river deposits on two different large fans. The Horikoshi gravel and mud Member of the Kasui Hills has a recurring environment of lake mud deposition based on the freshwater shell fossils produced, and a gravel layer near the river mouth. The Fukuroi Formation consists of similar fan deposits on the Iwatahara Plateau and Ogasa Hills, while the Kasui Hills are thought to be lake deposits derived from the Kasui Formation. The gravel beds on the Iwatahara Plateau and Kasui Hills contain quartz porphyry, which was supplied by the Tenryu River system, whereas the gravel beds on the Ogasa Hills lack such elements, suggesting that they were supplied by the Oi River system.

Depositional Processes and Palaeogeography

The Ogasa Group consists mainly of coarse-grained sediments formed ravelly fan-delta that started in Calabrian, which was formed when the Akaishi Mountains began to uplift on a large-scale and the Tenryu River and Oi River became the supply rivers of coarse-grained sediments [5]. According to Moriyama and Mitsuno [17], the Ina-Akaishi tectonic crastal movement, which uplifted a mountain range with a specific height of 2,500 m from Ina Valley to the Akaishi Mountains, started after about 1.8 Ma in the early Pleistocene. Suganuma et al. [18] suggest that full-scale uplift of the Akaishi Mountains began between 1.4 and 1.0 Ma, based on the contrast of widespread volcanic ash layers interbedded in fan deposits in the Ina Basin. In this study area, it is estimated that after about 1.8 Ma in the early Pleistocene after the deposition of the Kakegawa Group, the area was generally uplifted, the Iwatahara-Kasui area became a land area, the north-eastern to eastern foot of the Ogasa Hills area became shallow water and a fan-delta was formed between the mouth of the Tenryu River system to the coast. During the late depositional stage of the Soga Formation at 1.2 Ma, a submarine channel-filled gravel bed was formed at the northeastern foot of the Ogasa Hills from the Tenryu River system, and a full-scale fan-delta of coarse-grained sediment was formed in the eastern Ogasa Hills during the depositional stage of the Osuga Formation that started after 1.19 Ma, receiving gravel supply from the Oi River system on its northeastern side. This period roughly coincides with the start of the full-scale, rapid uplift of the Akaishi Mountainsrange. Based on the depositional ages and depositional environments of the various Formations of the Ogasa Group described so far, the depositional processes and palaeogeography of the Group are shown in the following four stages, A to D (Figure4).

Stage A (The Soga Formation: 1.78–1.19 Ma)

In the Iwatahara-Kasui area to the Ogasa Hills area facing the Pacific Ocean, there was a sea area where the continental slope spread offshore from the shoreface during the Soga Formation depositional period, and there was a fan-delta of the Tenryu River system and a coast continuous with it (Figure 4A). At the end of this period, at 1.2 Ma, there was an inflow of coarse-grained sediments from the Tenryu River system from the northwest direction, resulting in the formation of submarine channels and their burial (Figure 4A(a)).

Stage B (The Osuga Formation: 1.19–0.91 Ma)

Two depositional areas, the Ogasa Hills area and the nearby Iwatahara-Kasui area, resulted in contrastingly different depositional environments: fan-deltas and inner bays respectively. The formation of a large fan-delta of the Oi River, which juts offshore into the Ogasa Hills area, has resulted in the development of different depositional environments in the two neighbouring east-west areas. An inner bay or lagoon was formed in the Iwatahara-Kasui area from 1.07 Ma onwards (Figure 4B). In the Ogasa Hills area, a large fan-delta was formed at the beginning of the Osuga Formation depositional stage by the Oi River system flowing into the Takatenjin area (Figure 4B(a)), followed by the Oi River system flowing in from the northeast side of the Kakegawa. Subsequently, the fan-delta progressively advanced to the southwest, fan expanding to offshore (Figure 4B(b)). An inner bay or lagoon (Figure 4B(c)) was formed in the Iwatahara-Kasui area on its northwestern side, and mud, sand and gravel beds were deposited in the inner bay from the Tenryu River system. During this period, a large fan-delta is assumed to have formed in the lower reaches of the present-day Tenryu River (Figure 4B(d)), and the Iwatahara-Kasui area is estimated to have been a deep inner bay environment due to fan-deltas extending on both east and west sides.

Stage C (The Kasui Formation: 0.91–0.77 Ma)

The fan area extending offshore from the Iwatahara-Kasui and Ogasa areas was submerged by the transgression, but the inner part of the fan area was closed by offshore sandbars formed by wave erosion of the fan area, forming a freshwater lake. The lake then expanded to the north due to the transgression (Figure 4C).

Stage D (Fukuroi Formation: after 0.77 Ma)

Large fan-deltas were formed by the Tenryu River and the Oi River in both the Iwatahara Plateau and the Ogasa Hills, respectively, and fan areas spread out. Each of them moved southwards and buried the offshore continental slopes. However, a lake remained in the southern part of the Kasui Hills in this early stage (Figure 4D).

Figure 4: Depositional processes and paleogeographic evolution of the Ogasa Group. The dotted shapes in the background represent the current topography, and the solid lines indicate isobaths at 100 m intervals. The dotted areas represent depths shallower than 100 m. A, The depositional stage of the Soga Formation (1.78–1.19 Ma), (a) The Gansyoji submarine channel; B, The depositional stage of the Osuga Formation (1.19–0.91 Ma), (a) The Takatenjin delta plain, (b) Fan-delta of the Oi River during the depositional stage of the Ogasayama gravel to the Mitsutoge gravel Members, (c) The inner bay between the Oi and Tenryu River fans, (d) Fan-delta of the Tenryu River; C, During the depositional stage of the Kasui Formation (0.91–0.77 Ma), a lake appeared on the Ogasa and Kasui Hills; D, The depositional stage of the Fukuroi Formation (0.77 Ma~), The fans of the Tenryu and Oi Rivers developed to the south again.

Acknowledgment

The author would like to thank Ms. Kate Mariana, Geology, Managing Editor of Earth & Marine Sciences, for the publication of this paper.

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DOI: 10.31038/GEMS.2025741

Abstract

In this contribution, we demonstrate that samples from the Zinnwald greisen deposits also contain diamonds that crystallized under crustal conditions. The occurrence of diamond whiskers in quard and lonsdaleite in fluorite opens a wide field of sophisticated research.

Keywords

Raman spectroscopy, Supercritical fluids, Formation of diamonds in crustal regions, Diamond whiskers

Introduction

The study of some samples (quartz, cassiterite, zinnwaldite, fluorite) from the famous tin deposit of Zinnwald (Sn-W-Li deposit at the German/Czech border in the eastern Erzgebirge) has shown by the occurrence of diamond, diamond whiskers, and lonsdaleite, that these normally mantle-type minerals occur in relatively surface-near mineralization and are not uncommon [1-6]. Further studies on the Zinnwald mineralization, including the host granite, showed that the whole palette of carbon minerals is typical and present. In the past, the study of graphite and related minerals was ignored. As a result, many fundamental minerogenetic aspects could, therefore, not be correctly interpreted. The Formation of diamond and SiC whiskers (in quartz and beryl) clearly demonstrates the formation on the spot [1]. That means the formation occurs directly on the Earth’s crust, which is traditionally not possible. Furthermore, the here-described graphite or carbon differs from classic graphite (e.g., from Madagascar) by the extreme Raman band at low frequencies, which, according to Shea and Wall (2012) [7], are radial breathing modes (RBM). Such graphite nanotubes can include and transport a variety of components, including NaSn₂F₅, boron, and many others. The Formation of diamonds and Co. in the crustal range needs an acceptable explanation. Is the input of supercritical fluids and melts coming from the mantle range into the crust responsible for the exceptional mineralization? Is the transition from the supercritical stage to the critical and under-critical stages connected with pressure shock wave-like cavitation processes? Is the formation of spherical lonsdaleite crystals in fluorite with a diameter of about 40 µm or diamond whiskers in quartz with lengths up to 200 µm so feasible?. The following new observations on the Zinnwald samples, presented here, should help address these questions.

Sample Material and References

Granite Sample

The sample is a halogen-enriched topaz-albite granite from the Zinnwald tin-tungsten mining district (sample ZW212) of the Variscan age [8-10]. A detailed description of this and related samples and the results of fluid and melt inclusion studies can be found in Webster et al. (2004) [11] and Thomas et al. (2005) [12]. At that time, it was already clear that there was a two-stage formation of the granite, at least. The second phase exhibits a high fluorine content in the melt (~6% F). Topaz shows three different morphologic types: fine needles of topaz crystals, intergrowth of topaz, fluorite, and cryolite, as well as isometric topaz crystals [12]. Unconsidered up to now is the frequent appearance of graphite (Figure 1) and also nano- to micro diamonds in these granites. Striking is also the occurrence of spherical REE-rich fluorite crystals in quartz phenocrysts [13], which, together with the obtained pseudo-binary solvus curve (granite melt-water system) and the Lorentzian distribution of certain elements (F, Rb, Cs), demonstrate an input of supercritical fluids or melts from the Earths mantle into the crust [5,6].

Figure 1: Graphite (Gr) in granite from Zinnwald (sample ZW212) [12].

Pegmatite Quartz

Incorporated with the Variscan tin-tungsten mineralization related to the granites are pegmatite α-quartz crystals (up to 20 cm long), which are partially replaced by later β-quartz [1]. Other essential minerals are topaz, cassiterite, zinnwaldite, and fluorite. The last three contain, in part, diamond, diamond whiskers, lonsdaleite, graphite, graphite whiskers, and α-rhombohedral spheres and whiskers of boron partly with boron carbide [14].

Zinnwaldite and Fluorite

For the Zinnwald tin-tungsten deposit, Zinnwald-Cinovec is, according to Baumann et al. (2000), the 3-meter thick pegmatite of the type “Stockscheider” between quartz porphyry and albite granite characteristical. The pegmatite consists of quartz, potassium feldspar, and zinnwaldite. Some smoky quartz crystals up to up to 20 cm long can be found. In the zinnwaldite bundle, there are mostly violet to colorless fluorite crystals enclosed. Fluorite crystals contain lonsdaleite and diamond inclusions [4].

Microscopy and Raman Spectroscopy

We performed all microscopic and Raman spectroscopic studies with a petrographic polarization microscope (BX 43) with a rotating stage coupled with the EnSpectr Raman spectrometer R532 (Enhanced Spectrometry, Inc., Mountain View, CA, USA) in reflection and transmission. The Raman spectra were recorded in the spectral range of 0–4000 cm^-1 using an up-to-50 mW single-mode 532 nm laser, an entrance aperture of 20 µm, a holographic grating of 1800 g/mm, and spectral resolution ranging of 4 cm^-1. Generally, we used an objective lens with a magnification of 100x: the Olympus long-distance LMPLFLN100x objective (Olympus, Tokyo, Japan). The laser power on the sample is adjustable down to 0.02 mW. The Raman band positions were calibrated before and after each series of measurements using the Si band of a semiconductor-grade silicon single-crystal. The run-to-run repeatability of the line position (based on 20 measurements each) is ± 0.3 cm−1 for Si (520.4 ± 0.3 cm^-1) and 0.4 cm^-1 for diamond (1332.7 cm^-1 ± 0.4 cm^-1 over the range of 80–2000 cm^-1). The FWHM = 4.26 ± 0.42 cm^-1. FWHM is the Full-Width at Half Maximum. We also used a water-clear natural diamond crystal (Mining Academy Freiberg: 2453/37 from Brazil) as a diamond reference (for more information, see Thomas et al., 2023) [15]. The zero-point position for the Raman spectroscopic measurements is shown in Figure 2.

Figure 2: Position of the diamond line during the Raman measurements in the range from -40 to 2015 cm^-1 (sample 2453/37 from Brazil). The mean of 20 new measurements is 1332.2 ± 0.4 cm^-1.

Reference Material

Diamond and Graphite

As reference material for diamond and graphite, we used well- studied material. The diamond from the Mining Academy Freiberg, which we have used since the start of the Raman spectroscopic studies, is used daily to calibrate the device (see Figure 2). Additionally, for the same purpose, graphite from Madagascar is also used.

Since diamond, lonsdaleite, and graphite are in the central focus of this contribution, we provide the Raman reference spectra of diamond and graphite here (Figures 2 and 3).

Figure 3: Typical Raman spectrum of small plate-shaped graphite crystals from Madagascar.

Single-Walled Carbon Nanotubes (SWCNTs)

This pure material is from ProGraphite GmbH, Untergriesbach, Germany. We used this reference material because the Raman spectra show strong bands in the low Raman range, unlike graphite crystals from Madagascar. The following features [7] characterize the Raman spectrum of the used single-walled carbon nanotubes as reference: a strong G-band at 1570 cm^-1, a weaker D-band at 1332 cm^-1, a resonant 2D band at 2650 cm^-1, and between about 80 and 500 cm^-1 the radial breathing modes (RBMs). Figure 4 shows a typical Raman spectrum of single-walled carbon nanotubes (SWCNT).

Figure 4: Typical Raman spectrum of single-walled carbon nanotubes (SWCNTs).

Results

Topaz-Albite Granite from the Zinnwald

The granite contains rare cassiterite-brown microcrystals surrounded by a halo of colorless, primary probably β-quartz (see Figure 5). The determination of β-quartz is not simple because the surrounding quartz is α-quartz, and β-quartz transforms fast into α-form. The quartz of the hallo shows extremely high Raman lines at 126.0 and 204.7 cm^-1.

Figure 5: Diamond in granite quartz (sample ZW212) from Zinnwald. The diamond grain is 75 µm deep from the sample surface. The sample is 500 µm thick. D – diamond, α- Qtz – low-quartz, at the trapping probably β-Qtz – high-quartz. Under intense light, the diamond is in the center, cassiterite-brown, and at the rim, blue.

In Figure 5, a diamond crystal in granite quartz is depicted. Finding such crystals is not simple, as the granite sample contains many small, dark crystals, such as graphite. Figure 6 is a typical Raman spectrum of diamond shown – composed of the relatively broad diamond band at about 1326 cm^-1 and the typical G-band at 1585 cm^-1 of the carbon material. Depending on the isotope composition (¹³C vs. ¹²C), the band position of the diamond and the corresponding G-band can vary in a wide range [2]: pure ¹³C-diamond at 1283.1 cm^-1 and pure ¹²C-diamond at 1332.7 cm^-1. The same is the case for graphite: 1519 to 1581 cm^-1 for both end members (¹³C vs ¹²C) respectively.

Figure 6: Raman spectrum of diamond in granite from Zinnwald (sample ZW212, Figure 5).

The diamond is partially transformed into carbon (see Figure 7) because the diamond is at high temperatures and low pressure over a long time, not stable [16].

Figure 7: Carbon, transformed from primary diamond (Figure 6) by its ascent with supercritical fluids or melts into the present granite from Zinnwald (sample ZW212).

The direct diamond-related carbon is very rare in the granite sample ZW212. Most carbon is related to spheric or quasi-spheric black globules, often present and then mostly in quartz. Typical Raman spectra are in Figures 6 and 8. Characteristically, for this type of carbon, there are extreme Raman bands in the low-frequency range from about 80 to 400 cm^-1. The intensity is so strong that the range from 1200 to 2000 cm^-1 is nearby suppressed. Taking only the Raman spectrum in the range between 1200 and 2000 cm^-1 reveals the typical carbon band.The extremely high Raman bands are identified as radial breathing modes (RBM) by Shea and Wall (2012).

Figure 8: Raman spectra of graphite in granite from Zinnwald (ZW212), about 30 µm deep in quartz. Because the Raman band of graphite between 80 and 400 cm^-1 is exceptionally high (lower Figure), therefore the range of the upper Figure is strongly enlarged by using only the Raman window from about 1200 to 2000 cm^-1.

In contrast to crystalline graphite (see Figure 3) from Madagascar, the carbon in the Zinnwald granite shows extremely intense Raman bands in the low-frequency range. It looks like carbon nanotubes (SWCNTs) or MWCNTs (multi-walled carbon nanotubes – see Figure 4 and Shea and Wall (2012). The relatively strong band at 763.4 cm^-1 can traced back to fluoroboric acid H[BF₄] [12] (Figure 8).

Most carbon spheres in the granite quartz are composed obviously of multi-walled carbon nanotubes filled with different ions (rare alkalis, SnF₂, SnO₂, H₂O, and many see Figure 9).

Figure 9: Raman spectrum of SnF₂-loaded single-walled carbon nanotubes (SWCNTs). The RBM Raman band at 242.7 cm^-1 corresponds to a SWCNT diameter of about 1 nm filled with NaSn₂F₅, and the Raman bands at 483.3 and 621.2 results from SnO₂ [17].

Granite-related Tin-Tungsten Mineralization

Smoky Quartz

The results on smoky quartz are summarized in Thomas (2025a) [1]. Some carbon needles often have small diamond crystals at their tip (deep under the sample surface!). Figure 10a is a Raman spectrum of such carbon needles with a diamond on the tip.

Figure 10b show an about 80 µm long diamond whisker in quartz and the corresponding Raman spectrum. This whisker shows two points that held up. Chaotic processes during the crystal growth of the quartz host are so out of the question.

From the Raman band position, the diamond whisker is composed of 47.4% ¹³C, and the graphite is composed of 38.5% ¹³C, respectively.

Figure 10 a: Raman spectrum of a diamond-bearing carbon needle in quartz deep under the sample surface.

Figure 10 b: Diamond whisker in smoky quartz (upper photomicrograph) and the accompanying Raman spectrum of diamond whisker. The Raman band at 461.5 cm^-1 is from the quartz host.

Zinnwaldite

A brief description of the deposits in Zinnwald (Germany) and Cinovec (Czech Republic) is provided by Baumann et al. (2000) [8], Breiter and Qiao et al. (2024). The Li-mica zinnwaldite (~2.2 % Li₂O), named after this deposit, is the main constituent of this deposit and represents one of the largest Li-Sn-W greisen deposits in Europe. A careful study of zinnwaldite cleavage crystals reveals, besides small crystals of quartz, cristobalite, scheelite, cassiterite, and K₂SnCl₆, the presence of graphite grains. Figure 11 is an example of such a crystal with nanodiamonds (Table 1, Figures 11 and 12).

Table 1: Raman spectroscopic data of diamond and related carbon of the diamond in the Zinnwald granite ZW212; measurements on 19 crystals.

Sample	Diamond (cm-1)	FWHM	Carbon (cm-1)	FWHM
ZW212	1328.5 ± 4.3	63.1 ± 18.0	1578.7 ± 7.2	57.5 ± 13.5

Figure 11: Boron carbide-diamond crystal in zinnwaldite (Znw) from Zinnwald. The Raman spectrum of this crystal is in Figure 12.

Figure 12: Raman spectrum of boron carbide and maybe a diamond grain in zinnwaldite. The Raman band 1294 cm^-1 can be a hint to a ¹³C-rich diamond [18]. Exceptional are the strong 1448 and 1657 cm^-1 bands corresponding to boron carbide.

Besides such irregular grains, there are spherical crystals of a mixture of B_xC, BN, and Si₃N₄ with strong Raman bands between 400 and 1460 cm^-1 (see Figure 13). Similar spherical crystals are in smoky quartz [1] from Zinnwald and in cassiterite from Ehrenfriedersdorf (Sn-58) – Thomas, 1982. Boron carbides form at high temperatures and are stable at high pressure (GPa-range) and are also stable under intense Raman laser irradiation (50mW).

Figure 13: Raman spectrum of the spherical boron carbide B_xC crystal (shown at the top right) in zinnwaldite (Znw).

In the low-frequency range, a band appears at 187 cm^-1 (due to the B-C-B-C chain defect) as the pressure increases, as reported by Jay et al. (2023) [19]. After these authors, a significant kinetic barrier stabilizes the high-temperature configuration. Furthermore, in nature, it is probably a stoichiometric composition, not realistically – it is more a family of compounds of different compositions. However, the Raman spectra of such globules are more or less identical in zinnwaldite, quartz, and cassiterite. Very frequently, black films of graphite are observed. Figure 14 shows such aggregates.

Figure 14: Thick graphite (Gr) film in zinnwaldite (Zwn) from Zinnwald.

The Raman spectrum of such graphite is shown in Figure 15. Opposite to the graphite from Madagascar is FWHM (Full Width at Half Maximum), noticeably larger (see Table 2).

Figure 15: Raman spectrum of a graphite film in zinnwaldite (Figure 14), characterized by a strong Raman band at 1570.8 cm^-1. This value corresponds to approximately 16.5% ¹³C.

Table 2: Results of the Raman measurements on the graphite main peak (G band).

Sample	G band	FWHM	n (measured crystals)
Graphite in Zwn	1567.9 ± 4.6	21.8 ± 2.8	14
Gr from Madagascar	1580.2 ± 1.1	15.3 ± 1.9	14

Gr – graphite, Zwn – zinnwaldite.
Besides graphite, there are also inclusions filled with abiotic organic material (Figure 16).

Table 2 shows the Raman data for the graphite-G peak. In most cases, the D band, at approximately 1345 cm^-1, is mostly absent (Figure 16).

Figure 16 shows the inclusion of abiotic organic material in zinnwaldite. It is a complex mixture of saturated hexane (Author Collective, 1987 – [20]; Hurai et al., 2015 – [21]; Thomas, 2024d – [22]), produced by a natural Fischer-Tropsch process with zinnwaldite as the catalyst.

Figure 16: Complex abiotic organic material in zinnwaldite (Zwn) from Zinnwald.

Fluorite

In zinnwaldite, purple fluorite crystals (~5 mm in diameter) are often present. Thomas and Trinkler (2024) [4] found spherical lonsdaleite in these crystals. Besides the frequent lonsdaleite crystals, another rare diamond type is also present (Figure 17), characterized by two diamond Raman bands corresponding to high ¹³C contents.

Figure 17: Diamond in a fluorite crystal in zinnwaldite from Zinnwald/Saxony. The main band lies at 1313.7 ± 3.3 cm^-1 (10 measurements), which corresponds to 46.5% ¹³C.

Lonsdaleite is more frequently in the fluorite sample out of the zinnwaldite. Figure 18 shows the lonsdaleite crystal (a) and the Raman spectrum (b) from this crystal. In Table 3, the Raman-active vibrational modes are shown [23].

Figure 18: Lonsdaleite in fluorite from Zinnwald. a) Photomicrograph of lonsdaleite in fluorite. b) Raman spectrum of lonsdaleite. The Gaussian fit gives three components typical for lonsdalite.

Table 3: Raman-active vibrational modes.

Raman mode	Band position (cm-1)	FWHM (cm-1)
E2g	1245.6	36.1
A1g	1308.4	79.3
E1g	1356.5	24.4

According to Goryainov et al. (2018) [23], the three Raman-active vibration modes E_2g (here 1245.6 cm^-1), A_1g (1308.6 cm^-1), and E_1g (1356.5 cm^-1) are present.

The high ¹³C values of diamond in fluorite from Zinnwald spoke for a late supercritical input. The first input must be ¹²C richer.

Altenberg Tin Deposit

Not far from the Zinnwald deposit (~5 km) is the famous, however, closed Variscan tin deposit Altenberg [9,24]. Because there are also clear hints of supercritical fluids/melts present, the first data from that source should be given here in brief. Table 4 shows the results for diamonds found in quartz from pycnite rock. Pycnite is a variety of topaz, and the rock with the same name is composed of pycnite-topaz (44.2 ± 9.7), quartz (28.7 ± 8.8), and Li-mica (27.1 ± 8.0 %(vol/vol). The pycnite-topaz contains 19.8 % F [25].

Table 4: Raman spectroscopic data of diamond and related carbon of the diamond in the Altenberg pycnite rock; measurements on 18 crystals.

Sample	Diamond (cm-1)	FWHM	Carbon (cm-1)	FWHM	n
Pycnite-Qtz	1324.2 ± 10.2	74.8 ± 18.0	1585.8 ± 7.6	57.0 ± 7.8	18
Pycnite	–	–	1572.2 ± 7.1	20.8 ± 5.2	12

n – number of studied crystals.

The Altenberg example demonstrates that the pycnite cupola in the mine is the result of the interaction of a primary pegmatitic (quartz- feldspar) stock with an F-rich supercritical fluid into the pycnite rock. Remnants of potassium feldspar are present. Such interaction is also valid for the entire Greisen body.

Interpretation

In this contribution, further data are presented, which underline the interpretation of the interaction of supercritical fluid and/ or melts coming from the Earth’s mantle with the more crustal granites forming the granite-related Variscan tin mineralizations. The more or less subparallel with the granite cupola from Zinnwald [9] proceeding mineralizations contain diamond, graphite, boron carbide, and orthorhombic cassiterite, which are clear high-pressure and high-temperature indicators. Additionally, the formation of diamond whiskers in smoky quartz [1] clearly demonstrates that not all high-pressure minerals were transported with supercritical fluids. They crystalized on the spot. The finding of such minerals also in typical greisen formations (zinnwaldite) demonstrates the significant influence of supercritical fluids, which bring a lot of water, solid indicator minerals, and dissolved ore elements for the formation of greisen and pegmatite bodies. These scientific proofs necessitate a new interpretation of the formation of this type of Variscan tin mineralization. From our studies, it is clear that supercritical fluids or melts have a significant impact on ore and pegmatite formation. At the transition of the supercritical fluid into the critical and undercritical fluids at the formation site of the Variscan mineralizations, processes occur that are not yet well understood. The example of Zinnwald shows that such processes occur over a large spatial area, demonstrating that also the input of supercritical fluids is not only spatial restricted. Figure 19 shows the pressure-temperature diagram for the equilibria of diamond and graphite, and coesite and β-quartz.

Figure 19: Schematic phase diagram of diamond and graphite as well as between coesite and β-quartz. SCF – supercritical fluid, D – diamond, Gr – graphite, β-Qtz – high-quartz, Coe – coesite.

The transport of diamonds, as an example, with the supercritical fluid is not the problem. However, the formation of diamonds, for instance, as whiskers at a crustal level with about 0.1 to 0.3 GPa is, at the moment, a significant problem. One explanation is that during the sudden change from the supercritical state to the critical state, an enormous pressure rise occurs. Conceivable is, however, a more slow transition. Another possibility is that processes like cavitation have a significantly detrimental effect. However, cavitation processes are more chaotic. Here is already a vast playground of sophisticated research.

Discussion

In the last years, starting from 2023, the author and coauthors have found in minerals of Variscan tin deposits of the Erzgebirge (Germany), Krušné Hory (Czech Republic), and Slakoský les (Kaiserwald) in the Czech Republic high-pressure and high-temperature minerals in more crustal regions (pressure about 0.1 to 0.3 GPa. These minerals belong to diamond, lonsdaleite, moissanite, complex boron carbides, orthorhombic cassiterite, coesite, cristobalite-X-I, as well as beryl-II. Studied were the deposits Zinnwald, Altenberg, Cinovec, Krupka, Sadisdorf, Ehrenfriedersdorf, the granite from Greifenstein, granites from Annaberg, Schlaggenwald (Slavkovský les). Because samples can be contaminated during preparation [26-31] with diamond and SiC (moissanite), the samples were carefully cleaned, and only diamonds were used, which are far enough below the surface. The discovery of diamond- and moissanite whiskers in these samples enhanced the reliability of our findings. Furthermore, the proof of extremely ¹³C-rich diamonds and graphite in the studied samples clearly demonstrates the reality of our findings.

Acknowledgment

This contribution is dedicated to Adolf Rericha for his encouragement and his intense discussions of the problems related to the meaning of supercritical fluids.

References

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