Age and Origin of the Dongping Au-Te Deposit in the North China Craton Revisited: Evidence from Paragenesis, Geochemistry, and In Situ U-Pb Geochronology of Garnet

Gold vein deposits, including intrusion-related gold deposits (Sillitoe and Thompson, 1998; Lang and Baker, 2001) and metamorphic-related orogenic gold deposits (Goldfarb et al., 2005), account for a large proportion of global gold supply. Precise dating of gold vein deposits is needed to establish their temporal relationship to the geodynamic framework under which they formed and to develop a comprehensive genetic model. Numerous attempts have been made to constrain the timing of gold deposition, mainly by U-Pb dating of U-bearing accessory minerals, such as xenotime and monazite (Rasmussen et al., 2006; Fielding et al., 2017), ⁴⁰Ar/³⁹Ar dating of K-bearing alteration phases (Kent and McDougall, 1995; Li et al., 2003, 2012), and Re-Os or Rb-Sr isochron dating of gold-bearing sulfide minerals (Arne et al., 2001; Yang and Zhou, 2001).

Garnet contains variable U contents and has a high closure temperature of U-Pb isotopes (>850°C; Mezger et al., 1989), making it a potential geochronometer for geologic processes. The low diffusivity of U and Pb in garnet (Mezger et al., 1989) further helps to retain the good retentivity for the U-Pb isotope system. Recently, garnet has been utilized as a robust U-Pb chronometer to date igneous rocks and hydrothermal ore deposits (Deng et al., 2017; Seman et al., 2017; Gevedon et al., 2018; Li, D.F., et al., 2018; Wafforn et al., 2018; Burisch et al., 2019; Betsi et al., 2020; Duan et al., 2020). However, U-Pb dating of hydrothermal garnet remains in its infancy and has been limited to skarn genesis and associated mineralization. In this study, we test the applicability of garnet U-Pb dating to gold vein deposits, taking the Dongping Au-Te deposit as an example.

The Dongping Au-Te deposit (~120 t Au) is one of the largest gold deposits on the northern margin of the North China craton and has long been considered to be a Devonian alkaline-related gold deposit, largely based on the broad comparability of zircon U-Pb dates of high-grade orebodies (390–350 Ma; Li et al., 1998; Bao et al., 2014) to those of the host alkaline complex (400–390 Ma; Miao et al., 2002; Bao et al., 2014). However, Rb-Sr and ⁴⁰Ar/³⁹Ar dates of K-feldspar and sericite have bracketed gold mineralization to a younger, large age range of 289 to 103 Ma (Lu et al., 1993; Hu and Luo, 1994; Song and Zhao, 1996; Mo et al., 1997; Jiang and Nie, 2000; Hart et al., 2002; Cisse et al., 2018), which is broadly consistent with late Paleozoic to early Mesozoic orogenic deformation along the northern margin of the North China craton (Zhang et al., 2009; Zhai and Santosh, 2013). The younger age range together with structurally controlled veins, low sulfide contents in the veins, and low-salinity ore fluid has led to an orogenic gold model for Dongping (Hart et al., 2002; Cisse, 2016; Goldfarb et al., 2019). The lack of precise and reliable age constraints on the Dongping Au-Te deposit has hampered our understanding of how gold mineralization relates to the geodynamic evolution of the North China craton.

This study was prompted by our recent work on the Dongping Au-Te deposit, where we firstly identified hydrothermal garnets of various paragenetic stages. Herein, we integrated field and textural data, fluid inclusion microthermometric results, and geochemistry of garnets to document their hydrothermal origin. We then conducted laser ablation-inductively coupled plasma-mass spectrometry (LA-ICP-MS) U-Pb isotope analyses of garnet grains from each stage to date gold mineralization. To evaluate the relationship between gold veining and regional magmatism, magmatic garnet and/or zircon from the host syenite intrusion and a granite stock near the Dongping mine were also dated using the same technique. The geochronologic results, combined with field and geochemical data, are used to evaluate the timing and origin of the Dongping Au-Te deposit. The potential of garnet U-Pb dating as a robust geochronometer for gold vein deposits elsewhere is also discussed.

The North China craton is bounded by the Triassic Qinling-Dabie-Sulu orogenic belt to the south and the Paleozoic Central Asia orogenic belt to the north (Fig. 1A). It consists of the Eastern and Western blocks, which are separated by the ~300-km-wide Trans-North China orogen (Fig. 1A; Wilde et al., 2002; Zhao et al., 2005), which formed as a result of the collision between those two blocks at ca. 1.85 Ga (Zhao et al., 2005). From the late Paleoproterozoic to the end of the Paleozoic, the North China craton behaved as a stable cratonic block, characterized by weak magmatism and tectonism coupled with continuous sedimentation of marine carbonate and clastic rocks (Yang et al., 1986; Zhao et al., 2001). During the Mesozoic, the whole eastern half of the North China craton was tectonically reactivated, as marked by a large number of metamorphic core complexes or extensional domes (Fig. 1; Davis et al., 2002; Liu et al., 2005; Lin et al., 2011; Zhu, G., et al., 2015) and widespread rift basins containing voluminous bimodal volcanic rocks throughout the eastern North China craton (Ren et al., 2002; Zhu et al., 2010). Calc-alkaline magmatism of Early Cretaceous age was intense and widespread in the eastern North China craton (Fig. 1A; Wu et al., 2005). The above-mentioned extensional structures and large volume of igneous rocks are broadly coeval in the 140 to 120 Ma interval, and have been interpreted to result from significant thinning and destruction of the subcontinental mantle lithosphere keel triggered by the westward subduction of the Paleo-Pacific plate beneath the Eastern Asian continental margin (Griffin et al., 1998; Wu et al., 2005, 2019).

The Zhang-Xuan ore district on the northern margin of the North China craton is tectonically located in the Trans-North China orogen (Fig. 1A) and lithologically dominated by the Neoarchean Sanggan Group consisting of schist, gneiss, migmatite, and granulite (Fig. 1B). The schist has a whole-rock Rb-Sr isochron age of 2790 ± 155 Ma and a zircon U-Pb age of 2715 ± 21 Ma (Nie, 1998). The Sanggan Group is variably overlain by the Paleoproterozoic Hongqiyingzi and Changcheng groups (Fig. 1B). The Hongqiyingzi Group is distributed mostly to the north of the EW-trending Shangyi-Chongli-Chicheng fault and comprises marble, quartzite, amphibolite, and gneiss. The Changcheng Group is limited to the southeast of the Zhang-Xuan district and consists of a sequence of clastic marine sedimentary rocks (Nie, 1998). Both the Sanggan and Hongqiyingzi groups are locally covered by Early Cretaceous volcanic rocks of the Zhangjiakou Formation (Fig. 1B).

Large areas of the Zhang-Xuan district were affected by four major episodes of magmatism as represented by the Paleoproterozoic Wenquan granite (1709 ± 10 Ma; Nie, 1998), the Devonian Shuiquangou alkaline complex (400 ± 4–390 ± 6 Ma; Bao et al., 2014; Miao et al., 2002), the Triassic Guzuizi (236 ± 2 Ma; Miao et al., 2002) and Honghualiang granite (235 ± 2 Ma; Jiang et al., 2007), and the Early Cretaceous Shangshuiquan (143 ± 1 Ma; Jiang et al., 2007) and Beizaizi granite (131 ± 1 Ma; Miao et al., 2002) along with the Zhangjiakou Formation volcanic sequence (143 ± 3–136 ± 1 Ma; Wei et al., 2008) (Fig. 1B). The Shuiquangou alkaline complex intrudes the Neoarchean Sanggan Group (Fig. 1B) and is mainly composed of syenite and monzonite, but intrusive contacts between these lithologies are not obvious (Nie, 1998). The Guzuizi and Honghualiang granites are emplaced into the northern portion of the Shuiquangou alkaline complex or Neoarchean basement rocks (Fig. 1B), whereas the Early Cretaceous Shangshuiquan and Beizaizi granites occur to the south of the Shuiquangou alkaline complex and are generally overlain by volcanic rocks of the Zhangjiakou Formation (Fig. 1B). The Zhang-Xuan district is endowed with numerous Au-Te vein deposits that are hosted in the Shuiquangou alkaline complex, with Dongping being the largest one (Fig. 1B).

The Dongping Au-Te deposit is primarily hosted in the syenite intrusion of the Shuiquangou alkaline complex, with minor ores in surrounding metamorphic rocks of the Neoarchean Sanggan Group (Fig. 2). The deposit consists of ~70 orebodies that are structurally controlled by NNE-striking fracture zones and can be grouped into nine ore zones (Au-1 to Au-9; Fig. 2), with the Au-6 ore zone accounting for ~80% of the total gold resources. Individual orebodies consist of a number of quartz-sulfide veins that are commonly 45 to 600 m long and 0.2 to 10 m thick (Fig. 2; Bao et al., 2016). The ore grade varies mainly from 3 to 30 g/t with an average of ~10 g/t but locally reaches ~300 g/t (Nie, 1998). Most gold veins are enclosed by hydrothermal alteration envelope with an average width of 10 to 40 m that contains 1 to 5 g/t gold (Fig. 3).

The quartz-sulfide veins typically contain about 3 to 5 vol % sulfides dominated by pyrite and galena with minor amounts of chalcopyrite and sphalerite (Fig. 4A). Some auriferous quartz veins contain magnetite (Fig. 4B), whereas barren quartz veins usually have specularite (Fig. 4C). Gold-bearing pyrite and magnetite are also disseminated in the alteration envelopes surrounding the gold veins (Figs. 3, 4D, E). Free gold grains make up ~90% of the gold resources in the ores (Bao et al., 2016). The gold is typically intergrown with tellurides, chalcopyrite, and pyrrhotite (Fig. 5A, B) and occurs as mineral inclusions or microfracture infillings in pyrite, quartz, and locally in magnetite (Fig. 5C-E). Telluride minerals, dominated by calaverite, hessite, and petzite, are ubiquitous and commonly occur as mineral inclusions within pyrite (Fig. 5A, B) or locally within galena. Minor tellurides are also present in magnetite (Gao et al., 2015), but no gold or tellurides were observed in specularite. In general, gold grade is positively correlated with the abundance of tellurides (Gao et al., 2015).

Hydrothermal alteration is extensive in the Dongping Au-Te deposit. Potassic alteration is the most pervasive and largely expressed by thick K-feldspar veins (Fig. 6A), closely spaced veinlets (Figs. 4D, 6D) or wide K-feldspar halos (Fig. 6B-G). Flaky specularite aggregates usually occur in the potassic alteration zone (Figs. 4C, 6B, C). Minor garnet is present in K-feldspar-quartz veins and closely associated with specularite (Figs. 4C, 7A). Auriferous quartz-sulfide veins commonly dislocate or crosscut the preexisting K-feldspar veins (Fig. 6A, D), indicating that gold deposition postdates potassic alteration. In many cases, these quartz-sulfide veins and their alteration envelopes contain variable amounts of garnet and epidote that frequently overprint, replace, or crosscut the early K-feldspar and are closely associated with gold-bearing pyrite or magnetite (Figs. 6D-F, 7B-E). Some epidote grains are intergrown with native gold (Fig. 5C) and locally occur along microfractures of garnet (Fig. 7C). Sericite is also present in mineralized quartz-magnetite and quartz-sulfide veins (Figs. 4B, 6D, G). It generally replaces preexisting hydrothermal K-feldspar or garnet (Figs. 4B, 6D, G, 7E, F) and locally occurs as fine-grained aggregates in K-feldspar (Fig. 7F). Carbonate alteration is relatively weak and shows no relationship to gold deposition. Late quartz-calcite veins commonly crosscut the auriferous quartz-sulfide veins (Fig. 6H) and are deficient in sulfide (Figs. 6H, I, 7G). Minor amounts of garnet and epidote occur as cavity infillings in some quartz-calcite veins (Figs. 6I, 7G).

Based on field and paragenetic relations as mentioned above, three stages of hydrothermal alteration are recognized (Fig. 8): stage I K-feldspar, quartz, garnet, specularite, and minor pyrite; stage II quartz, garnet, epidote, sericite, magnetite, abundant sulfides, native gold, and telluride; and stage III quartz and calcite with minor garnet and epidote. These three stages of hydrothermal alteration are largely represented by pre-ore K-feldspar-quartz-specularite veins, syn-ore quartz-sulfide ± magnetite veins with sulfide-magnetite disseminations in associated alteration envelopes, and postore quartz-calcite veins, respectively.

A total of six samples were collected from the economically most important Au-6 ore zone (Figs. 2, 3) for this study. To test the possible relationship between Au-Te mineralization and regional magmatism, two magmatic rock samples each from the Shuiquangou alkaline complex and the Shangshuiquan granite close to the Dongping gold mine (Fig. 2) were selected for magmatic zircon and garnet U-Pb dating.

Samples DP-1, DP-8, and DP-10 were collected from the No. 70 orebody along the 1,264-m adit (Fig. 3). Sample DP-1 is from a pre-ore vein consisting of coarse-grained K-feldspar, quartz, and specularite, with minor amounts of light-brown garnet that is 0.1 to 1 mm across (Figs. 4C, 6B, C, 9A). Sample DP-8 is from a high-grade quartz-sulfide vein (Figs. 4A, 6E, 9B) containing ~100 g/t Au. This sample consists of quartz, pyrite, and garnet with minor magnetite and galena (Fig. 9B). Garnet with a 0.5- to 1-mm diameter in this sample is brown and honey yellow and closely associated with quartz, auriferous pyrite, and magnetite (Fig. 9B). Native gold and Au-bearing tellurides are abundant and typically hosted in pyrite. Sample DP-10 is taken from a post-ore quartz-calcite vein that intersects the ore-stage quartz-sulfide veins (Figs. 6H, 9C). It mainly consists of quartz, calcite, brown garnet, and epidote (Fig. 9C). Both euhedral garnet (0.2–1.5 mm across) and epidote commonly occur as cavity infillings in quartz-calcite veins (Fig. 9C).

Samples CK253-273, CK255-317, and CK256-359 represent mineralized alteration envelopes adjacent to different quartz-sulfide veins of No. 1 orebody beneath the No. 70 orebody and were taken from various drill holes (Fig. 3). These samples contain brown garnet grains (0.5–1.5 mm in diameter) and epidote that are spatially associated with disseminated gold-bearing pyrite and magnetite (Fig. 9D-F).

Sample DP-14 is an unaltered syenite from the Shuiquangou alkaline complex (Fig. 10A). This sample consists of K-feldspar (~50 vol %), albite (35 vol %), hornblende (10 vol %), and quartz (3–5 vol %), with accessory magmatic garnet, magnetite, apatite, titanite, and zircon (Fig. 10A). Sample DP-7 was taken from the Shangshuiquan granite and composed of orthoclase (~40 vol %), quartz (30 vol %), plagioclase (20 vol %), and biotite (5 vol %), with accessory zircon and magnetite (Fig. 10B).

All samples were prepared as standard double-polished thin sections, and their mineralogy and paragenetic sequences were determined using optical microscopy. Zircon grains from the syenite (DP-14) and granite samples (DP-7) were separated using conventional heavy liquid and magnetic methods and then handpicked under a binocular microscope. Garnet grains from the syenite and three vein samples (DP-1, DP-8, DP-10) were handpicked under a binocular microscope after crushing, cleaning in an ultrasonic bath with distilled water, and drying. The screened zircon and garnet grains were mounted in epoxy disks and polished to expose their interiors. Garnet grains in the alteration envelopes (CK253-273, CK255-317, and CK256-359) were analyzed in situ on the thin sections.

SEM and EPM analysis: Backscattered electron (BSE) imaging was used to characterize the morphological and textural features of garnets in each sample, using a FEI Quanta 200 environmental scanning electron microscope (SEM) and a MonoCL detector attached on a JXA-8100 electron microprobe (EMP) at the State Key Laboratory of Geological Processes and Mineral Resources (GPMR), China University of Geosciences, Wuhan. Major elements of garnet were analyzed using a JEOL JXA 8230 electron probe microanalyzer, equipped with wave and energy dispersive detectors at the Center for Material Research and Analysis, Wuhan University of Technology. The operating conditions consisted of an accelerating voltage 20 kV, a beam current of 20 nA, and a beam diameter of 5 μm. Natural silicate minerals were used as standards, including albite [Na], sanidine [K], chromite [Cr], rhodonite [Ca, Mn], hematite [Fe], pyrope [Al, Si], olivine [Mg], and rutile [Ti]. Spectral lines, peak time(s), offpeak background time(s), and average minimum detection limits (wt %) used for the wavelength dispersive spectrometry (WDS) analyses are as follows: K (Kα, 10, 5, 0.03), Na (Kα, 10, 5, 0.03), Cr (Kα, 10, 5, 0.04), Si (Kα, 10, 5, 0.03), Al (Kα, 10, 5, 0.025), Mg (Kα, 10, 5, 0.02), Ca (Kα, 10, 5, 0.018), Ti (Kα, 10, 5, 0.015), Mn (Kα, 10, 5, 0.04), and Fe (Kα, 20, 10, 0.035).

Fluid inclusion microthermometry: Fluid inclusions were studied to provide information on the conditions under which the garnets from quartz veins of each stage and the associated alteration envelopes formed. Double-polished thick sections of samples DP-1, DP-8, DP-10, and CK256-359 were prepared for fluid inclusion petrography and microthermometry. Microthermometric measurements were performed using a Linkam THMS 600 heating-freezing stage (–196° to 600°C) at the Fluid Inclusion Lab of the China University of Geosciences (Wuhan). The heating-freezing stage was calibrated at the melting points of CO₂ (–56.6°C), the ice-melting of H₂O (0.0°C), and critical point of pure H₂O (374.1°C) with synthetic fluid inclusion standards. Precision was estimated to be ±0.5°, ±0.2°, and ±2°C for the runs in the range of –120° to –70°C, –70° to 100°C, and 100° to 600°C, respectively. The phase transitions including the eutectic temperature, final ice-melting temperature (T_m.ice), and liquid-vapor homogenization temperature (T_h) were recorded. Fluid inclusions modified by postentrapment processes, such as necking and leakage, were excluded. Fluid inclusion assemblages (FIAs), which represent the most finely discriminated, petrographically associated group of inclusions (Goldstein and Reynolds, 1994), and some well-preserved isolated inclusions were chosen for microthermometric measurements. The total homogenization temperatures of fluid inclusions in each FIA with a variation of over 15°C were discarded, according to the rules proposed by Goldstein and Reynolds (1994). Given that the eutectic temperatures range from –22° to –20°C, the salinities (wt % NaCl equiv) of fluid inclusions were calculated from the final T_m.ice using the program HokieFlincs for the NaCl-H₂O system (Steele-MacInnis et al., 2012). The fluid density was also determined using HokieFlincs (Steele-MacInnis et al., 2012).

Trace element and U-Pb isotope analysis: Trace element and U-Pb isotope analysis of zircon and garnet were simultaneously conducted using an Agilent 7500a ICP-MS apparatus interfaced to a GeoLas 2005 laser ablation system with a DUV 193-nm ArF-excimer laser (MicroLas Germany) at the Wuhan Sample Solution Analytical Technology Co., Ltd. Detailed operating conditions for the laser ablation system and the ICP-MS instrument followed those described by Liu et al. (2008, 2010) and are briefly summarized here. The areas free of mineral inclusions and microfractures were selected for trace element and U-Pb isotope analysis. The laser beam diameter used was 32 μm for zircon, whereas a relatively large spot size of 60 μm was adopted for garnet to maximize count rates and optimize precision. A wire signal smoothing device is used for the laser ablation system, with which smooth signals can be produced even at a laser repetition rate of 1 Hz (Hu et al., 2012). A repetition rate of 5 Hz and an energy density of 6 J/cm² were applied to minimize elemental fractionation (El Korh, 2014). Argon was used as makeup gas and mixed with helium as the carrier gas via a T connector before entering the ICP. Nitrogen (flow rate of 2 mL/min) was added into the central gas flow (Ar + He) of the Ar plasma to decrease the detection limit and improve precision, consequently increasing the sensitivity for most elements by a factor of two to three (Hu et al., 2008). Each analysis incorporated a background acquisition of approximately 20 to 30 s (gas blank) followed by 50 s of data acquisition from the sample. No garnet U-Pb standards were available, thus laser-induced element fractionation and instrumental mass discrimination were corrected by normalization to the zircon standard 91500 (Wiedenbeck et al., 1995). Zircon standard GJ-1 was used as the secondary standard to monitor the precision and accuracy of the U-Pb dating results. Time-dependent drifts of U-Th-Pb isotope ratios were corrected using a linear interpolation (with time) for every five analyses according to the variations measured on zircon 91500 (Liu et al., 2010). Preferred U-Pb isotope ratios used for zircon 91500 are from Wiedenbeck et al. (1995). The final uncertainties were propagated from uncertainties of the preferred and measured 91500 values and the measured sample values (Liu et al., 2010). The obtained mean ²⁰⁶Pb/²³⁸U age for zircon GJ-1 is 602 ± 1 Ma (1σ; mean square of weighted deviates [MSWD] = 0.3; n = 32), which is consistent with the recommended value (601.95 ± 0.4 Ma; Horstwood et al., 2016). The glass NIST SRM 610 was used as the external standard to calibrate the trace element analyses of garnets, and ²⁹Si was used as an internal standard. The Si contents of the garnets were determined by EMP analysis. Off-line selection and integration of background and analyzed signals, time-drift correction, and quantitative calibration for U-Pb isotopes and trace element analyses were conducted using the ICPMSDataCal software (Liu et al., 2010).

Common Pb corrections were not applied due to the heterogeneous distribution and high variability of common Pb. Garnet U-Pb dates were determined using linear regressions on Tera-Wasserburg concordia (Tera and Wasserburg, 1972), and the lower intercept with the concordia curve is reported to represent the age of the garnets. All isotopic data for garnet and zircon grains were processed using Isoplot/Ex_ver.4.1 (Ludwig, 2010).

Trace element mapping of garnet: Trace element mapping was conducted to reveal the distribution of U and other trace elements in garnets, using a laser ablation system (Photon-Machines Analyte HE with 193-nm ArF) attached to a quadrupole-based ICP-MS (Agilent 7900) at the Ore Deposit and Exploration Centre (ODEC), School of Resources and Environmental Engineering, Hefei University of Technology. Detailed operating conditions, analytical procedures, and data processing followed those in Wang et al. (2017) and are briefly summarized here. The laser ablation system utilizes a two-volume ablation cell designed by Laurn Technic Pty. Ablation was performed in an atmosphere of UHP He (0.9 L/min), and upon exiting the cell, the aerosol is mixed with Ar (0.87 L/min) immediately after the ablation cell via a T connector before entering the ICP. The ICP-MS system was optimized daily to maximize sensitivity on isotopes of the mass range of interest, while production of molecular oxide species (i.e., ²³²Th¹⁶O/²³²Th) was kept less than 0.2%. Element maps were generated by ablating sets of parallel line rasters in a grid across the sample. A beam size of 20 μm and scan speeds of 20 μm/s were used in this study. A laser repetition of 10 Hz was chosen at a constant energy output of 80 mJ, resulting in an energy density of ~2 J/cm² at the target. A 20-s background was acquired at the start of scanning to allow cell washout, gas stabilization, and computer processing; a delay of 20 s was used after ablation. Standards NIST-610 and GSE-1G were analyzed at the start and end of each mapping for data calibration. Images were compiled and processed using the LaIcpMs software (Wang et al., 2017). Based on mean Si content of garnets from EMP analysis, ²⁹Si was used as an internal standard for calibration.

Garnets from pre-, syn-, and post-ore quartz veins and mineralized alteration envelopes show similar morphological and textural characteristics. These garnets are typically honey yellow to brown and variably associated with K-feldspar, quartz, specularite, magnetite, pyrite, gold, and calcite (Figs. 6E, G, 7A, D, G, 9A-F, 11A-G). They commonly display well-developed oscillatory zoning (Fig. 11A-D), although texturally homogeneous grains are also present locally (Fig. 11E). The cores of garnets from pre-ore K-feldspar-quartz-specularite veins are characterized by abundant porosities and extensive microfractures that are commonly filled with quartz and specularite (Fig. 11A). Garnets from ore-stage quartz-sulfide veins usually show a poikilitic texture with a large variety of hydrothermal mineral inclusions, such as quartz, epidote, galena, chalcopyrite, and magnetite (Fig. 11B, E-G). Native gold is present in garnet either as inclusions in other minerals (e.g., magnetite; Fig. 11F) or as isolated grains (Fig. 11G). Microfractures are locally developed within garnet grains and variably filled with sulfides and tellurides including chalcopyrite, calaverite, altaite, and petzite (Fig. 11H). Garnets from post-ore quartz-calcite veins commonly have fewer porosities in the cores and no mineral inclusions (Figs. 7G, 11C). Quartz, calcite, and epidote are main minerals associated with garnet, but sulfides are not observed (Figs. 6I, 7G, 9C, 11C). Garnets from mineralized alteration envelopes are closely related with gold-bearing pyrite (Fig. 11D, I). Microfractures are developed in many garnet grains and filled with quartz and calcite (Fig. 11D, I).

Magmatic garnets in the syenite are typically dark and occur as euhedral to subhedral grains with a diameter of 1 to 10 mm (Fig. 10A). They are texturally homogeneous and closely associated with various rock-forming minerals, such as K-feldspar and quartz (Figs. 10A, 12). In some cases, titanite, magnetite, and quartz occur as mineral inclusions within garnets (Fig. 12C).

Garnets from quartz veins of each stage and mineralized alteration envelopes contain abundant primary and pseudosecondary fluid inclusions (Roedder, 1984). Most fluid inclusions exhibit rectangular, oval, and negative crystal forms (Fig. 13A-G). Secondary fluid inclusions are also observed in some garnets, but they were not considered in subsequent measurements. Fluid inclusions occur as isolations (Fig. 13E, F) or form petrographically well-defined FIAs (Fig. 13A-D, G). Most FIAs are heterogeneously trapped with variable liquid/vapor ratios (Fig. 13A, B, D, G), but some homogeneous FIAs with similar liquid/vapor ratio are also present (Fig. 13C). Based on phase proportions at room temperature, fluid inclusions can be grouped as liquid- and vapor-rich types, which frequently coexist to form FIAs (Fig. 13A-D, G). Liquid-rich fluid inclusions are mostly 4 to 10 μm in size and have vapor bubbles that occupy 15 to 50 vol % of the inclusions (Fig. 13A-D, G). Vapor-rich fluid inclusions range from 5 to 20 μm in size, and the vapor bubbles usually account for ~80 to 95 vol % of the inclusions (Fig. 11A-G). Neither daughter crystals nor liquid CO₂ were observed in all fluid inclusions.

A total of 77 fluid inclusions from 21 FIAs and another 15 isolations were studied. All liquid-rich fluid inclusions homogenized into liquid phase, whereas vapor-rich varieties homogenized into vapor phase. The microthermometric results and calculated parameters are summarized in Appendix Table A1 and graphically illustrated in Figure 14.

Fluid inclusion assemblages in garnets from pre-ore K-feldspar-quartz-specularite veins (sample DP-1) have T_m.ice of –6.9° to –0.9°C and calculated salinities of 10.4 to 1.6 wt % NaCl equiv (App. Table A1; Fig. 14A). Total T_h of homogeneously trapped FIAs varies from 384° to 415°C and the minimum T_h from heterogeneously trapped FIAs is 375° to 396°C (Fig. 14A, B). Similarly, isolated fluid inclusions have T_m.ice of –5.2° to –2.4°C and calculated salinities of 4.0 to 8.1 wt % NaCl equiv (App. Table A1; Fig. 14A). They show T_h from 380° to 391°C (Fig. 14A). The calculated fluid densities range from 0.564 to 0.664 g/cm³ (App. Table A1).

Fluid inclusion assemblages in garnets from high-grade quartz-sulfide veins (sample DP-8) and mineralized alteration envelopes (sample CK256-359) yield T_m.ice of –7.0° to –0.9°C and calculated salinities of 1.6 to 10.5 wt % NaCl equiv (App. Table A1; Fig. 14A). Homogeneously trapped FIAs finally homogenized at the temperatures of 330° to 386°C, which are similar to the minimum T_h (330°–375°C) of heterogeneously trapped FIAs (Fig. 14A, B). Isolated fluid inclusions recorded T_m.ice values of –5.7° to –1.9°C, which correspond to calculated salinities of 3.2 to 8.8 wt % NaCl equiv (App. Table A1; Fig. 14). Their T_h values vary from 327° to 399°C (Fig. 14A). The fluid densities are calculated at 0.562 to 0.816 g/cm³ (App. Table A1).

Fluid inclusion assemblages in garnets from post-ore quartz-calcite veins (sample DP-10) show T_m.ice of –6.5° to –0.9°C and calculated salinities of 1.6 to 9.9 wt % NaCl equiv (App. Table A1; Fig. 14A). Total T_h of homogeneous FIAs ranges from 303° to 341°C and the minimum T_h of heterogeneous FIAs varies from 281° to 322°C (Fig. 14A, B). Isolated fluid inclusions have T_m.ice of –5.1° to –2.3°C and total Th of 287° to 361°C (App. Table A1; Fig. 14A). The calculated salinities based on the T_m.ice are 3.9 to 8.0 wt % NaCl equiv. The calculated fluid densities range from 0.628 to 0.816 g/cm³ (App. Table A1).

The results of EMP and LA-ICP-MS spot analyses of garnets are tabulated in Appendix Tables A2 and A3, respectively. The results show that garnets are grossular (Grs) to andradite (Adr) in composition (App. Table A2; Fig. 15).

Garnets from quartz veins of each stage and alteration envelopes consistently have large to moderate variations in compositions (App. Table A2; Fig. 15) that possibly reflect the oscillatory zoning as revealed in BSE images (Fig. 11A-D). The pre- and post-ore garnets (samples DP-1 and DP-10) range in composition from Adr₁₀₀Grs₀ to Adr₆₂Grs₃₅ with less than 6% of almandine, pyrope, spessartine, and uvarovite (App. Table A2; Fig. 15). They contain total rare earth elements (REEs) of 9 to 268 mg/kg and 35 to 364 mg/kg, respectively (App. Table A3), and are characterized by enrichment in light rare earth elements (LREEs) and depletion in heavy rare earth elements (HREEs), with strong to weak positive Eu anomalies (Eu/Eu* = 1.4–21.9 for sample DP-1 and 1.6–8.2 for sample DP-10; Fig. 16A, C). Garnets from ore-stage quartz-sulfide veins (sample DP-8) range in composition from Adr₁₀₀Grs₀ to Adr₁₉Grs₇₃, with 2 to 9% of almandine, pyrope, spessartine, and uvarovite (App. Table A2; Fig. 15). These garnets contain 3 to 72 mg/kg REEs (App. Table A3) and are generally characterized by LREE enrichment and HREE depletion, but the reverse is observed for some grains (Fig. 16B). Garnets from alteration envelopes (samples CK253-273, CK255-317, CK256-359) vary in composition from Adr₁₀₀Grs₀ to Adr₂₄Grs₆₉ (App. Table A2; Fig. 15) and have total REE composition of 2 to 205 mg/kg (App. Table A3). Most garnet grains from sample CK253-273 are enriched in LREEs and depleted in HREEs, but some have HREE-rich and LREE-poor patterns (Fig. 16D). Garnets from samples CK255-317 and CK256-359 have either LREE-rich and HREE-poor or HREE-rich and LREE-poor patterns (Fig. 16 E, F) and show strong to weak positive or weak negative Eu anomalies (Eu/Eu* = 0.6–22.7; App. Table A3; Fig. 16D-E).

Garnets from the syenite are relatively homogeneous in composition, ranging from Adr₈₅Grs₈ to Adr₇₇Grs₁₅ with minor (5–7%) almandine, pyrope, spessartine, and uvarovite (App. Table A2; Fig. 15). These garnets have relatively high total REEs (313–759 mg/kg; App. Table A3), and REE patterns are featured by HREE enrichment and LREE depletion with weak positive Eu anomalies (Eu/Eu* = 1.0–1.3; App. Table A3; Fig. 17).

In the LA-ICP-MS trace element maps, two garnet grains from ore-stage quartz-sulfide veins (sample DP-8) and postore quartz-calcite veins (sample DP-10) are compositionally zoned (Fig. 18). The abundance of U has a broad positive correlation with Al, Y, and REEs, but is reversely correlated with Fe (Fig. 18A, B). This observation indicates that U is mainly contained in the Al-rich zones of garnet grains.

U-Pb isotope data from garnets and zircons are tabulated in Appendix Table A4 and graphically illustrated in Figures 19 and 20. The sample descriptions and U-Pb dates are summarized in Table 1.

Garnet U-Pb dates: The error ellipses for laser spot analyses from seven garnet samples define well-constrained mixing lines (Figs. 19A-F, 20A) on Tera-Wasserburg diagrams, which are interpreted to be a mixture of common Pb and radiogenic Pb resulting from decay of U since garnet formation (Tera and Wasserburg, 1972). The spread of these data points allows the common Pb mixing line to be regressed to obtain a relatively precise lower intercept age.

Garnets from pre-ore K-feldspar-quartz-specularite veins contain 0.01 to 4.85 mg/kg Th and 1.53 to 10.38 mg/kg U, have Th/U ratios of less than 0.01 to 0.77 (App. Table A3), and yield a lower intercept date of 140 ± 6 Ma (MSWD = 2.1, n = 36; Fig. 19A). Garnets from high-grade quartz-sulfide veins contain 0.01 to 0.39 mg/kg Th and 0.92 to 13.89 mg/kg U, corresponding to Th/U ratios of less than 0.01 to 0.25 (App. Table A3), and show a lower intercept date of 140 ± 4 Ma (MSWD = 1.5, n = 47; Fig. 19B). Garnets from post-ore quartz-calcite veins contain 0.01 to 1.07 mg/kg Th and 1.14 to 10.16 mg/kg U and have Th/U ratios of less than 0.01 to 0.22 (App. Table A3). The lower intercept date is constrained at 139 ± 6 Ma (MSWD = 1.3, n = 16; Fig. 19C).

Garnets from alteration envelopes have U-Pb dates that are comparable to those of various quartz veins (Table 1). These garnets contain 0.07 to 5.75 mg/kg Th and 0.80 to 11.69 mg/kg U with Th/U ratios from less than 0.01 to 0.78 (App. Table A3). Garnets from samples CK253-273, CK255-317, and CK256-359 have lower intercept dates of 142 ± 5 (MSWD = 1.4, n = 15), 141 ± 4 (MSWD = 0.7, n = 26), and 141 ± 6 (MSWD = 1.7, n = 29), respectively (Fig. 19D-F).

In comparison, garnets from the syenite (sample DP-14) of the Shuiquangou alkaline complex contain 0.11 to 5.52 mg/kg Th and 2.30 to 11.39 mg/kg U and have Th/U ratios of 0.05 to 0.50 (App. Table A3). They have a lower intercept date of 395 ± 6 Ma (MSWD = 1.2, n = 11; Fig. 20A).

Zircon U-Pb dates: Zircon grains from the syenite (sample DP-14) are mostly subhedral to euhedral prisms of 50 to 200 μm in length, with elongation ratios of 1 to 3. They usually display a well-developed oscillatory zoning in cathodoluminescence (CL) images, indicating a magmatic origin. These zircon grains have concordant or slightly discordant ²⁰⁶Pb/²³⁸U dates ranging from 399 to 383 Ma (App. Table A4), with a weighted mean of 392 ± 2 Ma (MSWD = 1.3, n = 17; Fig. 20B).

Zircon grains from the Shangshuiquan granite (sample DP-7) are colorless and subhedral to euhedral prisms with length of 50 to 150 μm and elongation ratios of 1 to 2. In CL images, these zircons are generally dark, with a few grains that exhibit oscillatory zoning typical of magmatic zircons. The ²⁰⁶Pb/²³⁸U dates of these oscillatory zoned zircons range from 147 ± 2 to 141 ± 2 Ma (App. Table A4) and form a coherent group on the concordia, with a weighted mean of 143 ± 1 Ma (MSWD = 1.3, n = 16; Fig. 20C).

The heterogeneous FIAs are petrographically well defined in hydrothermal garnets (Fig. 13A, B, D, G), and coexisting liquid- and vapor-rich fluid inclusions in these FIAs show similar homogenization temperatures (App. Table A1; Fig. 14A), indicating fluid boiling during garnet formation. Thus, the temperature and pressure conditions at which fluid inclusions were trapped can be directly constrained using these heterogeneously trapped FIAs (Goldstein and Reynolds, 1994). The minimum homogenization temperatures of heterogeneous FIAs represent their formation temperatures (Diamond, 2003; Bauer et al., 2019). The microthermometric results reveal a fluid cooling trend from 396° to 375°C (mean 383°C) in the pre-ore stage through 375° to 330°C (mean 348°C) in the ore stage to 322° to 281°C (mean 302°C) in the post-ore stage (Fig. 14A, B). The fluid entrapment pressures estimated for pre-, syn-, and post-ore stages are ~260 to 330, 70 to 160, and 90 to 145 bar, respectively (Fig. 14B), according to the boiling liquid-vapor curve for NaCl-H₂O system (Driesner and Heinrich, 2007).

Garnets of various stages are cogenetic with a sequence of ore and gangue minerals (Figs. 4C, 6E, I, 7A, C-E, G, 11A-D, I), such as K-feldspar, quartz, pyrite, and calcite, and contain abundant primary fluid inclusions (Fig. 13A-G), suggesting their precipitation directly from hydrothermal fluids. This view is further confirmed by the large variation in REE compositions and patterns (Fig. 16A-F), which are distinct from those of magmatic garnets from the syenite (Fig. 17) and likely reflect changes in fluid pressure, composition, and/or growth rates (Jamtveit et al., 1993; Ciobanu and Cook, 2004; Gaspar et al., 2008). Moreover, the flat time-resolved depth profile for U and Th (Fig. 21) indicates lattice-bound occurrence of U in garnet, a view partly supported by the absence of U-rich minerals (Fig. 11A-I) and the positive correlation between U and REEs (Fig. 18A, B). The U-Pb isotopes of hydrothermal garnets studied, therefore, record the time elapsed since their formation (Deng et al., 2017; Seman et al., 2017). Microthermometric results of fluid inclusions reveal that the formation temperatures of garnets of various stages (302°–383°C; Fig. 14A, B) are significantly lower than the closure temperature of U-Pb isotopes in garnet (>850°C; Mezger et al., 1989), further confirming that the hydrothermal garnet U-Pb dates can be considered as their crystallization ages. Because garnets from high-grade quartz-sulfide veins and mineralized alteration envelopes precipitated coevally with gold-bearing sulfides (Figs. 6E, 7D, 9B, D-F, 11D-I), their U-Pb dates can be also interpreted as the age of the Dongping Au-Te deposit. Garnet U-Pb dates of syn-ore quartz-sulfide veins are within errors of those of mineralized alteration envelopes (Fig. 19B, D-F), attesting to their quality and reliability when interpreted as the timing of gold deposition. Although garnet U-Pb dates of pre-, syn-, and post-ore quartz veins are indistinguishable within analytical uncertainties (Fig. 19A-C), it is unlikely to bracket the duration of hydrothermal processes and gold mineralization due to the large errors of the age data. Collectively, our new garnet U-Pb dates demonstrate the Dongping Au-Te deposit formed at ca. 140 Ma in the Early Cretaceous.

Previous studies attempted to constrain the formation age of the Dongping Au-Te deposit using zircon U-Pb isotopes, but no consensus has been reached. The marginal consistency of zircon U-Pb dates of high-grade quartz veins (390–350 Ma; Li et al., 1998; Bao et al., 2014) with those of the host syenite (400–390 Ma; Miao et al., 2002) has led to a prevalent view that the Dongping Au-Te deposit is genetically associated with the Shuiquangou alkaline complex and thus can be classified as an alkaline-related gold deposit (Zhang and Mao, 1995; Li et al., 1998; Mao et al., 2002; Nie et al., 2004; Cook et al., 2009; Bao et al., 2016). However, other quartz veins and alteration envelopes have zircon U-Pb dates ranging from 242 ± 7 to 140 ± 1 Ma (Bao et al., 2014; Li, H., et al., 2018). It is noteworthy that unequivocal textural and compositional evidence that can document the origin and geologic significance of zircon grains (Yeats et al., 1996; Schneider et al., 2012; Deng et al., 2015) was not presented or was insufficiently presented in previous U-Pb geochronology studies for Dongping. We speculate that those zircon grains are most likely fluid-entrained detrital zircons from the ore-hosting fracture zones that crosscut the syenite (Figs. 2, 3, 4A). The large U-Pb age range (400–140 Ma) suggests that those detrital zircons have been subjected to different degrees of hydrothermal alteration during ore-forming processes, leading to variable, partial loss of radiogenic Pb. The above considerations suggest that zircons extracted from quartz veins are mostly inherited grains from the host rocks, rather than precipitated directly from the ore fluids, and thus may not be a suitable phase for dating gold mineralization at Dongping.

The significant difference between our new garnet U-Pb dates and previous radiometric dating results for the ore-hosting syenite and gold veins suggests that both the Devonian alkaline-related (Zhang and Mao, 1995; Li et al., 1998; Mao et al., 2002; Nie et al., 2004; Cook et al., 2009; Bao et al., 2016) and orogenic models (Hart et al., 2002; Cisse, 2016; Goldfarb et al., 2019) can be precluded. Critical diagnostic features for orogenic gold deposits, including intense carbonate alteration, CO₂-rich fluid inclusions, and arsenic-rich pyrite, are generally lacking or atypical in the Dongping Au-Te deposit, as evidenced by our geologic observations (Fig. 6), fluid inclusion petrography (Fig. 13), and previous pyrite geochemistry study (Cook et al., 2009). In addition, no orogeny-related regional metamorphism has been documented in the interior of the North China craton since this craton was finally stabilized in the late early Paleoproterozoic (Zhao et al., 2005). More importantly, the Dongping Au-Te deposit formed in an extensional geodynamic framework (see below), which is different from the compressional to transpressional regimes typical of orogenic gold deposits. Taken together, we argue against an orogenic model for the Dongping Au-Te deposit.

Gold mineralization at Dongping (ca. 140 Ma; Fig. 19) is broadly coeval with widespread granitoid magmatism over the eastern North China craton (Fig. 1A; 152–130 Ma; Wu et al., 2005; Zhang et al., 2014), such as the Shangshuiquan granite (ca. 143 Ma; Fig. 20C). Many of granitoid intrusions contain abundant mafic microgranular enclaves and are associated with coeval mafic dikes, stocks, and volcanic rocks (Chen et al., 2003; Wang et al., 2011; Zhang et al., 2014). The abundance of mafic dikes and enclaves indicates lithospheric extension during the Late Jurassic to Early Cretaceous, which is further manifested by the development of metamorphic core complexes (Fig. 1A; 150–120 Ma; Guo et al., 2012; Wang et al., 2012) and rift basins in the eastern North China craton (145–110 Ma; Ren et al., 2002; Zhang et al., 2005). Similarly, most gold vein deposits in the eastern North China craton formed at 140 to 120 Ma and are spatially associated with metamorphic core complexes and mafic to intermediate dikes (Li et al., 2003, 2012; Tan et al., 2012; Zhu, R.X., et al., 2015). Therefore, the Dongping Au-Te deposit, as is the case for many other equivalents over the eastern North China craton (Fig. 1A), was emplaced under an extensional setting and is genetically related to the Early Cretaceous magmatism, which resulted from the thinning and destruction of the lithospheric keel beneath the North China craton, a geodynamic process induced by the westward subduction of the Paleo-Pacific plate (Wu et al., 2005, 2019).

A magmatic origin of the Dongping Au-Te deposit is supported by the magmatic-like H-O isotope signature of ore fluids (δ¹⁸O_fluid = 1.5–6.9‰, δD = –108 to –73‰; Nie, 1998; Fan et al., 2001) and mantle-derived He isotopes in the ore fluids (³He/⁴He= 0.3–5.2 Ra; Mao et al., 2003). Sulfur isotope data (δ³⁴S = –13 to 1.7‰; Song and Zhao, 1996; Nie, 1998) provide additional evidence for a magmatic source of sulfur and, by inference, other fluid components. Negative δ³⁴S values likely reflect isotopic fractionation driven by the oxidation of magmatic-derived fluids (Ohmoto, 1972; see below).

Mineralization and alteration features of the Dongping Au-Te deposit, such as extensive potassic alteration, predominance of hydrothermal andradite, common presence of iron oxides, and abundant tellurides (Figs. 4–6, 9), are broadly comparable with those of oxidized intrusion-related gold deposits (Robert, 2001; Helt et al., 2014; Mériaud and Jébrak, 2017). The common presence of specularite and magnetite in the pre- and syn-ore veins (Figs. 4B, C, E, 5D, 9A, B, E) indicates that the ore fluids were relatively oxidized, likely exsolved from oxidized magmas (Thompson et al., 1985; Jensen and Barton, 2000; Robert,, 2001; Helt et al., 2014). Fluid boiling as a critical mechanism for gold deposition in intrusion-related gold deposits (Jensen and Barton, 2000; Helt et al., 2014) is common at Dongping, as indicated by the occurrence of heterogeneous FIAs with similar homogenization temperatures (Figs. 13, 14). Although the Dongping Au-Te deposit is coeval with the nearby Shangshuiquan granite (Figs. 1B, 2, 19, 20C), it is highly unlikely that it was the source of ore fluids, because it is barren and lacks high-temperature alteration. We speculate that ore-forming fluids were sourced from an unexposed intrusion beneath the Dongping Au-Te deposit, presumably an equivalent to the Shangshuiquan granite (Figs. 1, 2).

The age of gold vein deposits, as best represented by intrusion-related and orogenic gold deposits, must be correctly determined to identify the relationship of gold mineralization to geodynamic framework (Sillitoe and Thompson, 1998; Groves et al., 2003). Garnet, commonly dominated by grandite, has been reported as a minor alteration phase in many intrusion-related gold deposits (App. Table A5), such as the Chalice gold deposit in the Yilgarn craton (Bucci et al., 2002) and several deposits in the Tombstone gold belt of Canada (Baker and Lang, 2001). Our study showed that hydrothermal grandite in the intrusion-related gold deposits can contain parts-per-million-level, lattice-bound U, which is sufficient to generate precise U-Pb dates. Garnets have also long been recognized in many orogenic gold deposits where they are closely associated with ore and gangue minerals (App. Table A6). These garnets are also dominated by grandite, but almandine, grossular, and spessartine are not uncommon in several deposits from the Appalachian orogenic belt in Canada (App. Table A6). Information on the U-Th budgets of grandite from orogenic gold deposits is essentially lacking, but one recent analysis on almandine from the Mupane gold deposit in Botswana revealed a sub-parts-per-million level of U and yielded a reliable U-Pb isotope date (Betsi et al., 2020). Thus, the common presence of U-bearing garnet closely associated with native gold in many intrusion-related and orogenic gold deposits (App. Table A5, A6) makes garnet a potential robust U-Pb geochronometer for gold mineralization.

In the Dongping Au-Te deposit, hydrothermal garnets are recognized in each vein stage and associated alteration envelopes, which provides an ideal opportunity to date gold mineralization. These garnets contain abundant boiling fluid inclusions and formed at the temperature of 302° to 383°C and estimated pressure of 90 to 330 bar. They have 0.80 to 13.89 mg/kg U residing predominantly in the crystallographic sites and are generally characterized by LREE enrichment and HREE depletion with prominent Eu anomalies. LA-ICP-MS spot analyses of garnets of various paragenetic stages yield indistinguishable U-Pb dates of 142 ± 5 to 139 ± 6 Ma (1σ), which constrain the formation of the Dongping Au-Te deposit at ca. 140 Ma in the Early Cretaceous, rather than in the Devonian or early Mesozoic as previously suggested. Combined with independent geologic and geochemical studies, results presented here suggest that the Dongping Au-Te deposit can be considered to be an oxidized intrusion-related gold deposit and was emplaced in an extensional setting related to the thinning and destruction of the lithospheric keel beneath the North China craton. An additional implication of this study is that garnet may serve as a potential robust U-Pb geochronometer for gold vein deposits.

This research was financially supported by the Ministry of Science and Technology of China (2016YFC0600104), the 111 Project (B20045), the Fundamental Research Funds for the Central Universities, China University of Geosciences, Wuhan (CUGCJ1711), and the GPMR State Key Laboratory (MSF-GPMR03). Fangyue Wang is thanked for providing access to the LA-ICP-MS mapping analysis facility. We also thank Yabin Zhang, Wei Huang, and Yuanhai Wang from the Dongping gold mine for their help with field work. Profs. John W. Valley and Philip E. Brown read an early version of this paper and are sincerely thanked. Thorough and constructive comments by two anonymous reviewers and Massimo Chiaradia, associate editor of Economic Geology, have been very helpful in our revision and are gratefully acknowledged. Our sincere thanks go to Prof. Larry Meinert, editor of Economic Geology, for his effective handling and constructive suggestions. This is a contribution 31 from China University of Geosciences Center for Research in Economic Geology and Exploration Targeting (CREGET).

Gao-Hua Fan received his bachelor’s degree in economic geology from China University of Geosciences (Wuhan) in 2016 and is currently a Ph.D. student at the same university. His B.Sc. thesis is on the origin of a Cu-Au skarn deposit in eastern China using mineral characterization, isotopic dating, and fluid inclusions. He currently works on lode gold deposits in the North China craton with an aim to better understand the timing, source, and evolution of gold mineralization in the context of craton lithospheric destruction by integrating structural analysis, drill core logging, mineral imaging, isotopic analysis, and geochemical modeling.