Probing amagmatic fluids in porphyry deposits via a mono- intrusion system at Tongchang (Dexing, SE China)

doi:10.21203/rs.3.rs-2726797/v1

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Probing amagmatic fluids in porphyry deposits via a mono- intrusion system at Tongchang (Dexing, SE China)

https://doi.org/10.21203/rs.3.rs-2726797/v1

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published 31 Jan, 2024

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Porphyry deposits contain large amounts of base metals, and bear important fingerprints of magmatic and hydrothermal processes in convergent and collided plate margins. For decades, source of amagmatic fluid and its role in sulfide deposition in porphyry systems has been equivocal.

The Tongchang mono-intrusion porphyry deposit is studied for fluid sources and depositional mechanisms. In-situ quartz O and anhydrite Sr isotope analyses revealed a coexistence of magmatic and amagmatic (groundwater and residue metamorphic) fluid reservoirs, spatially separated by an impermeable zone formed by quartz deposition near brittle-ductile transition in phyllite. The impermeable zone sealed inside amagmatic components. The early and early-intermediate fluid stages were dominated by magmatic fluids forming high-T disseminated Cu sulfides (>400 °C). The late-intermediate stage was marked by retreating of the brittle-ductile interface and creation of an intermediate zone below the impermeable quartz seal. Rupturing of the intermediate zone at 350 °C to 300 °C caused incursion of the sealed amagmatic fluids leading to deposition of vein-type Cu sulfides. The late fluid stage was characterized by massive invasion of evolved groundwater and ore remobilization in response to volumetric contraction of the fluid system. Four universal models are envisaged to account for variabilities in amagmatic availability and thermal histories in porphyry deposits.

Earth and environmental sciences/Solid earth sciences/Geochemistry

Earth and environmental sciences/Solid earth sciences/Geology/Economic geology

SEM-CL

anhydrite

O isotope

Sr isotope

fluid-rock interaction

Porphyry deposits are the world’s largest sink of copper and molybdenum, and a major source of gold, silver, and rhenium¹. They form via physico-chemical interaction of hydrous magma with surrounding rock/fluid reservoirs in the crust^2,3,4,5. For decades, fluid inclusion and isotopic studies have demonstrated the presence of amagmatic components (groundwater, connate water, metamorphic water, etc.) in porphyry deposits^6,7,8,9,10. Nevertheless, type of amagmatic components, timing and mechanism of incursion, and its relation to sulfide deposition remains hotly debated. Orthomagmatic models insist on late incursion of amagmatic components as cooling agents^9,10,11,12. Magmatic fluids are considered as the predominant and occasionally sole player in ore formation^{13,14,15,16,17,18}. In contrast, recent studies have discovered early-stage amagmatic incursion and coeval ore formation, which challenges the original orthomagmatic views^19,20,21,22. A good example is the Bingham Canyon porphyry Cu-Au-Mo deposit, in which chalcopyrite was recently found to precipitate from a hybrid fluid containing appreciable amounts of meteoric waters²⁰. These new discoveries make it imperative to re-assess the existing fluid models and pinpoint the role of amagmatic components for porphyry deposits.

Mono-intrusion porphyry systems are of particular interest because they may have simpler fluid history compared to multi-intrusion ones^23,24. In this study, the mono-intrusion Tongchang porphyry deposit is studied with oxygen and strontium isotope analyses. Previous O isotope studies for this deposit utilized bulk-rock method and, not unexpectedly, ended up with conflicting models. Spatial variability of O isotopes in the altered rocks suggest a main role of amagmatic fluids^25,26. By contrast, oxygen isotope analyses on quartz and phyllosilicate minerals support orthomagmatic models^27,28. This discrepancy, common in porphyry deposits, are primarily due to incapability of bulk analysis in detecting sub-mineral isotopic heterogeneity¹⁹. Additionally, O isotope systematics readily change with fluid-rock interaction and thus are not good indicator of the original fluid source^6,9,10. Frequently, radiogenic Sr isotopes have been used as a complementary tool ^7,29. Therefore, this study combines in-situ O isotope analyses of quartz and Sr isotope analyses of anhydrite for the Tongchang porphyry Cu deposit with the main objective of pinpointing amagmatic components.

Detailed descriptions of the regional and deposit geology can be found in several publications ^30,31,32. For the sake of simplicity, only the most relevant aspects are described in this paper.

The Tongchang porphyry Cu deposit along with the Fujiawu and Zhushahong deposits are located in a Neoproterozoic orogen that sutured the Yangtze and Cathaysia blocks ³¹. The three deposits, unlike most others, are associated with single porphyry intrusion. The three intrusions, with granodiorite composition, emanate from a common batholith at depths ³³ (Fig. 1a). The granodioritic magmas are derived from partial melting of a subduction-modified lithosphere at ca. 170 Ma ^30,31 (Fig. 1a).

Types and spatial patterns of hydrothermal alteration, vein and mineralization is broadly comparable to the typical ones ⁴. Potassic and propylitic alterations developed in the granodiorite and wallrock (mainly phyllite), respectively, and define concentric zones around the granodiorite stock. These early-formed alteration zones were overprinted sequentially by chlorite-sericite and sericitic alterations (Fig. 1b). Drill core logging suggests that there is a continuous zone of strong silicification in the phyllite (core column in the Fig. 1).

Hydrothermal veins at Tongchang are comparable to the A-B-D classification. Three A-type veins, two B-type veins, and two D-type veins are used in this study: hematite-quartz veins (abbreviated as A.V_hq, “A” denotes A-type, “h” and “q” the subscripts represent hematite and quartz, arranged in an increasing volume. The same rule applies to all other vein type nomenclatures), K-feldspar-quartz veins (A.V_kq), anhydrite-quartz veins (A.V_aq), chalcopyrite-pyrite-quartz veins (B.V_cpq), molybdenite-bearing anhydrite-chalcopyrite-pyrite-quartz veins (B.V_macpq), molybdenite-bearing quartz-pyrite veins (D.V_mqp), and pyrite-dominant veins (D.V_pd). Minerology and crosscutting relation suggest that A.V_hq, A.V_kq, A.V_aq are broadly comparable to the so-called A-type veins, B.V_cpq, B.V_macpq, to the B-type veins, and D.V_mqp and D.V_pd. to the D-type veins. Copper minerals, mainly chalcopyrite and subordinately bornite and Cu-sulfosalts, are localized in the chlorite - sericite and sericitic zones (Fig. 1b). In the chlorite-sericite zone, chalcopyrite occurs in veinlets (B.V_cpq and B.V_macpq) and disseminations in the altered host rocks. In the sericitic alteration zones, chalcopyrite occurs as mineral inclusions, and tennantite as disseminations in the altered rocks ³⁵.

SEM-CL has revealed that quartz veins at Tongchang are mostly composite, and formed via reopening of early quartz veins ³². Six generations of quartz have been recognized, and their formation temperatures have been estimated with fluid inclusion and TitaniQ geothermometry ³² (Table 1). The earliest quartz (Qz1) is characterized by mottled-CL texture (Fig. 2a), and they were formed at temperatures between 600 ºC and 650 ºC. The second quartz generation (Qz2) is characterized by an oscillatory CL core (Qz2a) and CL-mottled rim (Qz2b), and was formed at temperatures between 529 ºC and 655 ºC (Qz2a), and between 473 ºC and 584 ºC (Qz2b). The third quartz generation (Qz3) commonly contains a CL-dark interior (Qz3a) and oscillatory overgrowth (Qz3b) (Fig. 2b), and was formed at a temperature range of 374 ºC ‒ 420 ºC (Qz3a), and 350 ‒ 450 ºC (Qz3b). The fourth quartz generation (Qz4) is low in abundance and shows dark CL (Fig. 2a), and was formed at temperatures 310 ‒ 390 ºC. The fifth generation of quartz (Qz5) commonly consists of a CL-dark core (Qz5a) and oscillatory overgrow (Qz5b) (Figs. 3a, 3b), and it was formed at temperatures between 356 ºC and 451 ºC (for Qz5a) and between 346 ºC and 439 ºC (for Qz5b). The latest quartz generation (Qz6) consists of CL-dark to grey homogeneous domain (Qz6a) and very bright-CL domain (Qz6b) (Fig. 3a). The Qz6 quartz were formed at temperatures between 288 ºC and 370 ºC (Qz6a) and between 241 ºC and 313 ºC (Qz6b).

The first generation of anhydrite (Anh1) exhibits bright homogeneous to slightly patchy CL. They are subdivided to Anh1a and Anh1b according to CL intensity (Fig. 4a). The second generation (Anh2) has bright to grey CL with “wavy” oscillatory zoning (Fig. 4a). Anh2 is subdivided to CL-brighter Anh2a and CL-darker Anh2b. The third generation (Anh3) has euhedral morphology and display slightly oscillatory or homogeneous CL (Fig. 4a). The fourth generation (Anh4) is anhedral and CL-dark (Fig. 4b).

Textural evidence suggests the following paragenesis with quartz: Qz1 with hematite; Qz2a with Anh1a; Qz2b with Anh1b; Qz3a with Anh2a; Qz3b with Anh2b and pyrite; Qz4 with Anh4, chalcopyrite, pyrite; and Qz5 with pyrite. Quartz microtexture and deposition temperature allows dividing the Tongchang fluid history into early (Qz1 and Qz2), early - intermediate (Qz3), late - intermediate (Qz4), and late (Qz5 andQz6) stages.

3.1 Oxygen isotopic composition of quartz and fluid

Oxygen isotope results are listed in the Table 2 and depicted in Figs. 2, 3, 5. Qz1 has clustered δ¹⁸O values of 9.0‰ ‒ 9.3‰ (av. 9.2‰, N = 6). Qz2 has slightly larger δ¹⁸O range of 8.8‰ ‒ 10.2‰ (av. 9.2‰, N = 20) for Qz2a, and 8.1‰ ‒ 10.4‰ (av. 9.2‰, N = 44) for Qz2b. Qz3b has higher δ¹⁸O values of 10.2‰ ‒ 13.0‰ (av. 10.9‰, N = 27). Qz4 has distinctively high δ¹⁸O of 14.1‰ ‒ 19.6‰ (av. 16.6‰, N = 3); Qz5 has δ¹⁸O range of 11.2‰ ‒ 12.9‰ (av. 12.0‰, N = 8) for Qz5a and 10.4‰ ‒ 13.6‰ (av. 11.7‰, N = 17) for Qz5b. Qz6 peaks in δ¹⁸O at 13.9‰ ‒ 20.4‰ (av. 16.5‰, N = 9) for Qz6a and 16.6‰ ‒ 27.2‰ (av. 23.9‰, N = 16) for Qz6b. Above all, several low δ¹⁸O values (4.4 ‒ 8.9‰) are detected in a relictic quartz in a phyllite-hosted D.V_pd vein (Fig. 3b).

Oxygen isotope modeling suggests that hydrothermal fluids responsible for early and early - intermediate stages (Qz1, Qz2, and Qz3) have δ¹⁸O values of 7‰ to 8‰ (Fig. 5). The late-intermediate stage fluid (Qz4) has elevated δ¹⁸O (11 ‒ 12‰). Qz5 - precipitating fluid has δ¹⁸O values similar to magmatic water and evolved groundwater (7 ‒ 9‰). Qz6-depositing fluid has high yet variable δ¹⁸O values (9.4 ‒ 11.8‰ for Qz6a and 14.2 ‒ 17.1‰ for Qz6b).

3.2 Strontium isotopic composition of anhydrite and fluid

Strontium isotope results are listed in the Table 3 and depicted in Fig. 4, 6a. Anh1, Anh2, and Anh3 have ⁸⁷Sr/⁸⁶Sr ratios of 0.70457 ‒ 0.70527. A slight decreasing trend could be seen from Anh1 (0.70513) and Anh2 (0.70509) to Anh3 (0.70478). By contrast, Anh4 has evidently higher ⁸⁷Sr/⁸⁶Sr ratios (0.70583 ‒ 0.70646). ⁸⁷Rb/⁸⁶Sr ratio was less than 0.001 for all analyzed spots, and therefore, radiogenic Sr is considered negligible. Therefore, the measured ratios are considered representatives of the anhydrite-depositing fluids.

4.1 Possible fluid sources

There are three possible types of parental fluids in the Tongchang deposit. Presence of magmatic fluids was demonstrated by previous H-O isotope data ^27–28. The magmatic fluids had δ¹⁸O of 7.0 ‒ 8.8‰ and ⁸⁷Sr/⁸⁶Sr of 0.70455 ³³. Groundwater (δ¹⁸O: -8.2‰) was suggested by whole-rock O isotope analyses ^25–26. In-situ O isotope analysis of this study revealed a ¹⁸O-depleted quartz grain (δ¹⁸O: 4.4‰, Fig. 3b), likely indicating equilibrium with this groundwater at ca. 200 ºC. The host terrain was metamorphosed at greenschist conditions ³¹, and metamorphic fluids may have been preserved. Metamorphic fluids in the adjacent Jinshan orogenic gold deposit had δ¹⁸O values between 10.4‰ and 11.2‰ ³⁷.

The isotopic composition of these fluids could change via rock interaction. Mass balance calculation suggests that groundwater could become ¹⁸O-enriched depending on temperature and w/r ratio (δ¹⁸O up to 13.0‰ at 600 ºC and w/r 0.001, Fig. 6b). Meanwhile, deuterium become either enriched at < 400 ºC or depleted at > 400 ºC. Despite many successful applications ⁶, hydrogen isotope method is not always straightforward due to fluid overprints. In addition to O and H isotopes, Sr isotopes of the groundwater could approach that of the phyllite during rock interaction. An age-corrected ⁸⁷Sr/⁸⁶Sr_(170Ma) of 0.71215 was reported for phyllite ²⁵, which is taken as the Sr isotope ratio of the evolved groundwater.

Similar w/r reactions may occur for the metamorphic fluid, especially considering its long residence time between regional metamorphism (ca. 800Ma) and ore formation (ca. 170Ma). At > 300 ºC, the evolved metamorphic fluids become ¹⁸O-enriched and D-depleted, and vice versa for temperature < 300 ºC (Fig. 6c). δ¹⁸O value up to 13.0‰ can be attained at 600 ºC and w/r 0.01.

4.2 Pinpointing amagmatic components

Paragenetic relation and O-Sr isotope analyses suggest that early and early - intermediate fluids (Qz1, Qz2-Anh1, Qz3-Anh2, and Anh3) may have had a common fluid source. Magmatic water, evolved groundwater, and metamorphic fluid are possible candidates based on O isotopes (Fig. 5).

⁸⁷Sr/⁸⁶Sr ratios of the early anhydrites (Anh1-3: 0.70457 ‒ 0.70527) requires mixing between magmatic fluid and amagmatic components. The relative proportion depends on Sr content (denoted as [Sr] thereafter) of the two fluids. 10 ppm [Sr] is assumed for the amagmatic fluids based an overview of global metamorphic fluids by ³⁸. 10 ppm [Sr] is also used for the magmatic fluids based on studies of the Butte and Bajo de la Alumbrera porphyry deposits ^39,40. Modeling suggests a predominance of magmatic component for Anh1, Anh2 and Anh3 (92%, 93% and 97%, respectively), equivalently, 3–8% amagmatic components (curve 1 in the Fig. 6a). Higher amagmatic [Sr] values decrease the amagmatic proportion (curve 2 in the Fig. 6a), whereas higher magmatic [Sr] values up to 100 ppm increase the amagmatic proportion to ca. 45% (curves 3 in the Fig. 6a).

For late-intermediate stage fluid (Qz4-Anh4), using [Sr] of a similar magmatic fluid of the Butte porphyry deposit (ca. 480 ppm ³⁹) and 100 ppm [Sr] for amagmatic ⁴¹, calculation suggests 43% of magmatic and 57% of amagmatic components (curve 4 in the Fig. 6a). This estimation, if correct, requires highly-evolved groundwater and/or metamorphic water (δ¹⁸O: 12‰) as the amagmatic source. This is quite likely since these fluids may have existed at Tongchang as elaborated in the previous section. However, the low w/r ratio (< 0.01) required for the w/r reaction seems contradictory to the large volume of Qz4-Anh4 veins and associated alteration. This contradiction may be reconciled by a model for the Bingham porphyry Cu deposit⁶. The authors envisioned an exchange zone where small patches of groundwater became ¹⁸O-enrich at high temperature and low w/r, and subsequently migrated and accumulated at the ore-deposition zone.

Fluid-rock interaction models for equilibrium fractionation between groundwater (G.W.) (b), with metamorphic water (M.W.) (c), and phyllite and Rayleigh distillation (d). H-O isotope values of groundwater are from ²⁶; Metamorphic water values are from ³⁷. In (d), the initial composition is 8‰ and the incremental vapor loss is 10% per 10°C from 340°C to 250°C.

δ¹⁸O values of the Qz5 - precipitating fluids are generally indistinguishable from magmatic water and evolved groundwater (Fig. 5). Two pieces of evidence support a dominant role of amagmatic fluid. First, bulk-rock Sr isotope analyses suggest that sericitic altered porphyry had large ⁸⁷Sr/⁸⁶Sr ratios (up to 0.709), indicating dominance of phyllite-derived Sr (curve 5 in the Fig. 6a). Second, the large volume of sericitic alteration requires large volumes of fluids that cannot be provided only by a cooling magma chamber ⁴².

Another prominent feature for the late stage is the high yet variable δ¹⁸O values in Qz6 with some being the highest for porphyry deposits (up to 27‰). The parent fluids also have high and variable δ¹⁸O values (from 9.4‰ to 17.1‰). The large δ¹⁸O variation can neither be explained by fluid cooling nor fluid/rock interaction. It is worth noting that the lowest fluid δ¹⁸O values are close to the evolved groundwater and metamorphic water prevailed in Qz5. If the same fluid existed at the beginning of Qz6, a Rayleigh distillation process can explain the observed high and variable δ¹⁸O. It requires a gradual cooling from 340 ºC to 250 ºC and vapor loss up to 70% (Fig. 6d).

4.3 Mechanisms of amagmatic fluid incursion

Amagmatic proportions decreased from 8–3% in the early to early-intermediate stages. These estimates are not trivial because the difference of ⁸⁷Sr/⁸⁶Sr of Anh1 to Anh3 are much larger than analytical precision. The decreasing trend may reflect a close-system behavior, where available amagmatic components are gradually consumed. Mechanism of forming a closed system in early porphyry environment was vigorously simulated by ⁴³. The author demonstrated formation of an impermeable zone through retrograde precipitation of quartz and the brittle to ductile transition (BDT). The impermeable zone separates the internal magmatic fluids from the external groundwater.

Nevertheless, the model ⁴³ does not consider possibility of trapping amagmatic components within the impermeable zone, which is essential for explaining the high δ¹⁸O amagmatic signatures. Additionally, the model considers only multi-intrusion systems, where heats were continuously supplied by new pulses of magma, as opposed to a progressively cooling mono-intrusion system of Tongchang. A mono-intrusion system would see downward retreating of the BDT, and formation of an intermediate region between the two interfaces. This intermediate region is broadly similar to what is called “carapace” in orthomagmatic models ⁴⁴. It is under a hydrostatic regime, and is intermittently ruptured due to overpressure from the underlying cupola.

The late-intermediate stage at Tongchang, may have seen a major rupturing of the intermediate region, which triggered the releasing of the amagmatic components trapped during the formation of the impermeable zone (Fig. 7). Meanwhile, the quartz shell may have remained intact or less ruptured. A relatively intact quartz shell would prevent a massive inflow of external groundwater. Even though it is not entirely clear why quartz shell remained intact, one explanation would be its farther distance from the cupola. Subsequently at the late stage, magmatic fluid production may have been significantly decreased such that the fluid pressure inside the magmatic - hydrothermal region was greatly reduced, causing volumetric contraction and breach of the silica shell. The breach of the silica shell induced large invasion of evolved groundwaters with δ¹⁸O values similar to those of magmatic fluids.

4.4 Role of amagmatic fluid incursion in Cu deposition

During the early to early-intermediate stages at Tongchang, Cu sulfides deposited mainly as disseminations in the potassic and chlorite ‒ sericite altered rocks, a common phenomenon in porphyry deposits ⁴⁵. These early sulfides may have deposited from magmatic fluids at high temperatures (650 ºC to 400 ºC) via chemisorption reaction ⁴⁶.

Different from most porphyry deposits, vein-type chalcopyrite deposition at Tongchang coincided with the incursion of amagmatic waters. This temporal coincidence likely suggests amagmatic incursion as a promoting factor for Cu deposition based on the several pieces of argument. First, the early-intermediate stage experienced less intense Cu deposition with less amagmatic components, even though other physico-chemical conditions (fluid acidity, composition, and pressure) were similar. Secondly, the amagmatic fluids are more dilute and cooler than magmatic fluids, the inflow of which has been demonstrated to enhance of Cu sulfide deposition ¹⁷ due to dramatical destabilization of ${CuCl}_{2}^{-}$ complexes and elevation of ${H}_{2}S$ activity ^43,47. Lastly, the amagmatic fluids buffered at oxidizing conditions are capable of remobilizing Cu sulfides through partial dissolution ⁴⁸.

A high to intermediate sulfidation assemblage (pyrite and tennantite) developed in the phyllic altered rocks, which reflects stability of muscovite, quartz, and pyrite in the expense of chlorite, feldspars, mafic minerals, and chalcopyrite. This is similar to the Butte porphyry deposit where the deep protores (chalcopyrite-pyrite-magnetite) with potassic-chloritic alterations were remobilized by circulating oxidized groundwaters ⁴⁹.

4.5 Establishing and extending the Tongchang model

A genetic model is proposed for the Tongchang porphyry deposit (Fig. 7). It commences with the establishment of magmatic cupola, which is followed by the formation and evolution of the silica shell and BDT.

The early and early-intermediate stages are dominated by magmatic fluids, which produced low-sulfidation assemblages in the potassic zone (Fig. 7a), and cooled to generate intermediate sulfidation assemblage in chlorite-sericite zone (Fig. 7b). The late-intermediate stage is marked by accumulation of amagmatic fluids (evolved groundwater and metamorphic fluid), subsequently mixing with magmatic fluids, and massive deposition of chalcopyrite in veins (Fig. 7c). The late stage, characterized by wanning of the magmatic input and dominance of evolved groundwater, was responsible for remobilization of preexisting metal sulfides (Fig. 7d). The Tongchang model should be applicable porphyry Cu deposits that comprise mono-intrusion surrounded by amagmatic fluid reservoirs.

Figure 7A genetic model for the formation of silica shell and brittle/ductile transition in the Tongchang porphyry Cu deposit. (a) In the early stage, retrograde precipitation of quartz formed a silica shell, coinciding with the brittle/ductile transition (BDT). Amagmatic components were trapped within the ductile phyllite. Disseminated chalcopyrite was formed by rock reactions of magmatic fluid with mafic minerals. Early barren quartz veins were formed in fractures. (b) In the early-intermediate stage, the BDT moved downward, leaving behind a hydrostatic intermediate zone, which was ruptured upon fluid overpressures and formation of intermediate-T disseminated ores. (c) In the late-intermediate stage, the trapped amagmatic components migrated upward and mixed with magmatic fluids, forming vein-type Cu sulfides. (d) Magmatic input vanned and induced massive invasion of evolved groundwater and remobilization of earlier ores.

The Tongchang model may be extended to account for variability in the amagmatic availability and intrusive history (Fig. 8a). In porphyry deposits without amagmatic fluids, such as those emplaced in contemporaneous volcanic and igneous rocks ⁵⁰, thick silica shell could not be formed. In such system, fluid evolution would be dominated by magmatic fluids and thus resemble the early and early-intermediate stages of the Tongchang model and typical orthomagmatic models ⁵ (Figs. 8b, c). In this scenario, ores would be disseminated sulfides associated with potassic alteration in the case of multiple intrusions (Fig. 8b), or with chlorite-sericite alteration in the case of single intrusion ⁵⁰ (Fig. 8c).

In multi-stock systems involving amagmatic fluids, common in most porphyry deposits ⁴, three scenarios could be distinguished depending on the spatial configuration of the intrusion, which affects the thermal history of the system. If repeated magma injection maintains the brittle-ductile transition at a location close to the silica shell until mineralizing fluid is consumed (Fig. 8d), little to limited amounts of amagmatic fluid would be available in the intermediate zone, so ores would be dominated by high-T disseminated sulfides, which would be remobilized by external groundwater at low temperatures. In contrast, if repeated magma injection occurs in such a way to allow incursion of trapped amagmatic components (Fig. 8e), ore styles and spatial distribution would resemble that of Tongchang. An example would be the El Salvador deposit 51 (Gustafson and Hunt, 1975), where the pre-mineral stocks act merely as country rocks. If repeated magma injection causes the silica seal to breach at high temperatures, magmatic fluids would surge out and displace external groundwater circulation (Fig. 8f). This scenario, similar to the one illustrated by ⁴³, would result in phreatic brecciation of the country rocks, allow breccia-related porphyry-type mineralization at depth and epithermal mineralization atop.

Combined oxygen and Sr isotope analyses provide new evidence for fluid processes at the magmatic-dominant to amagmatic dominant transition in the Tongchang mono-intrusion porphyry deposit. Possible amagmatic components included chemically and isotopically evolved groundwater and metamorphic residual fluids. The hydrological evolution was largely controlled by interplay between an impermeable shell formed by retrograde solubility and BDT. The initial amagmatic incursion coincided with the major period of Cu sulfide deposition. The diluting and cooling effects on magmatic fluids are interpreted as a promoting factor for Cu mineralization. The Tongchang model is applicable to mono-intrusion porphyry deposits where amagmatic components are available. The Tongchang model is extended to three scenarios that have more widely implications.

Nine selected quartz vein samples cover all vein and alteration types. Guided by SEM-CL petrography, quartz is analyzed for oxygen isotope with secondary ion mass spectrometry (SIMS) and anhydrite is analyzed for Sr isotopes with multi-collector LA-ICP-MS. Water-rock interaction and mixing models are constructed based on O-Sr isotope data.

6.1 In-situ O isotope analyses

In-situ quartz oxygen isotopic compositions were measured using a Cameca IMS 1280 SIMS at the Institute of Geology and Geophysics, Chinese Academy of Sciences (IGGCAS). A Cs⁺ primary ion beam with an intensity of about 2 nA was accelerated at 10 kV and focused on an elliptic spot of ~ 10 × 20 µm in size. Oxygen isotopes were measured using multi-collection mode on two off-axis Faraday cups, with ¹⁶O counts being typically 109 cps during the analytical session. Obtained ¹⁸O/¹⁶O ratios were normalized to the VSMOW waters (¹⁸O/¹⁶O: 0.0020052), and then corrected for instrumental mass fractionation using NBS-28 (δ¹⁸O_VSMOW: 9.5‰ ⁵²). A quartz standard (Qinghu, δ¹⁸O: 8.4 ± 0.2‰) was used to monitor instrumental drift. Measured δ¹⁸O values for Qinghu are between 8.3 and 9.1‰, with a mean of 8.6 ± 0.2‰ (1σ, n = 22). Analytical uncertainty for each analysis was normally better than 0.3‰ ‒ 0.4‰ (2σ).

6.2 In-situ Sr isotopes

Strontium isotopes in anhydrite were determined with a Neptune Plus MC-ICP-MS equipped with a MICRO/Las Pro 193 nm ArF excimer laser ablation system at IGGCAS. Laser parameters were set at a repetition rate of 8 Hz, a fluence of 10 J/cm², and 90-µm spot size. Ion signals of ⁸³Kr, ⁸⁴Sr, ⁸⁵Rb, ⁸⁶Sr, ⁸⁷Sr, and ⁸⁸Sr were monitored by six Faraday cups at low-resolution mode. Prior to analysis, the ICP-MS was tuned to a maximum sensitivity using a standard solution. Analytical runs started with measurement of Kr gas for 40 seconds with laser-off, followed by laser ablation for 60 seconds. A piece of shark tooth (⁸⁷Sr/⁸⁶Sr: 0.709179 ± 0.000007, 2σ) as the primary standard, and Slyudyanka apatite (⁸⁷Sr/⁸⁶Sr: 0.70766‒0.70773) as the quality-check standard, were analyzed twice every ten unknowns ⁵³. Data reduction was conducted offline and potential isobaric interferences were resolved following the protocols described in ⁵³. Results of the Slyudyanka apatite suggested an analytical precision of ± 0.000113 (1σ) and accuracy of ± 0.000129.

6.3 Oxygen and Sr isotope modeling

Oxygen isotopic compositions of fluids in equilibrium with quartz were calculated with the following equations ⁵²:

${{\delta }^{18}O}_{qz}-{{\delta }^{18}O}_{fluid}=2.05\times \left({\left(T+273.15\right)}^{2}\right)-1.14$ (500ºC <T< 800ºC)

${{\delta }^{18}O}_{qz}-{{\delta }^{18}O}_{fluid}=3.34\times \left({10}^{6}/{\left(T+273.15\right)}^{2}\right)-3.31$ (250ºC < T<500ºC)

Water-rock interaction models were derived using the following mass-balance Eq. 5⁴:

$${C}_{w}\times w\times {HO}_{H2O}^{i}+{C}_{r}\times r\times {HO}_{r}^{i}={C}_{w}\times w\times {HO}_{H2O}^{f}+{C}_{r}\times r\times {HO}_{r}^{f}$$

where $C$, $w$, $i$, $r$, $f$, and $HO$ stands for wt.% H or O, water, initial value, phyllite, final value, and δD or δ¹⁸O.

The H and O fractionation factors for phyllite (${\delta }^{18}{O}_{r}^{f}-{\delta }^{18}{O}_{H2O}^{i}$; $\delta {D}_{r}^{f}-\delta {D}_{H2O}^{f}$) were approximated by muscovite $2.68\times {10}^{6}\times {T \left(in K\right)}^{-2}-3.57$ ⁵⁵, and $-22.1\times {10}^{6}\times {T \left(in K\right)}^{-2}+19.1$ ⁵⁶, respectively.

Modeling of Sr mixing is based on the following Eq. 5⁷:

$${f}_{b}=\frac{{\left(\frac{{}^{87}Sr}{{}^{86}Sr}\right)}_{a}-{\left(\frac{{}^{87}Sr}{{}^{86}Sr}\right)}_{m}}{{\left[{}^{86}Sr\right]}_{b}/{\left[{}^{86}Sr\right]}_{a}\times \left({\left(\frac{{}^{87}Sr}{{}^{86}Sr}\right)}_{m}-{\left(\frac{{}^{87}Sr}{{}^{86}Sr}\right)}_{b}\right)+{\left(\frac{{}^{87}Sr}{{}^{86}Sr}\right)}_{a}-{\left(\frac{{}^{87}Sr}{{}^{86}Sr}\right)}_{m}}$$

where ƒ_b denotes weight percentage of amagmatic fluid; subscripts m, a and b represent anhydrite, magmatic fluid, and amagmatic fluid; [⁸⁶Sr] denotes ⁸⁶Sr concentration.

Acknowledgements

We thank Yueheng Yang and Qiuli Li for their kind guidance during MC-LA-ICP-MS and SIMS analyses. X.L. especially acknowledges Hongrui Fan for his unwavering help with sample and data collection for the last few years. This study was collectively supported by the Extra & Co project (Valorisation-Instituts Carnot n°ANR-15-CNRT-003), the National Natural Science Foundation of China (41402083), and the Traceability of Primary Mineral Resources and Recovery and Recycling of Secondary Mineral Resources (P22017).

Author contributions

X.L. and K.Y. initiated the idea and work plan; X.L. performed the experiments; X.L., A.R., J.P., and K.Y. interpreted the data; X.L. prepared the manuscript; A.R., J.P., and K.Y. edited the manuscript. X.L. and J.P. acquired research funds.

Data availability statement

All data generated or analyzed during this study are included in this published article.

Competing Interests Statement

The authors declare no competing interests.

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Probing amagmatic fluids in porphyry deposits via a mono- intrusion system at Tongchang (Dexing, SE China)

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Abstract

Figures

1. Introduction

2. The Tongchang Deposit

3. Results

3.1 Oxygen isotopic composition of quartz and fluid

3.2 Strontium isotopic composition of anhydrite and fluid

4. Discussion

4.1 Possible fluid sources

4.2 Pinpointing amagmatic components

4.3 Mechanisms of amagmatic fluid incursion

4.4 Role of amagmatic fluid incursion in Cu deposition

4.5 Establishing and extending the Tongchang model

5. Conclusion

6. Methods

6.1 In-situ O isotope analyses

6.2 In-situ Sr isotopes

6.3 Oxygen and Sr isotope modeling

Declarations

Acknowledgements

Author contributions

Data availability statement

Competing Interests Statement

References

Tables

Additional Declarations

Supplementary Files

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