The relationship between base metal deposits, especially Mississippi Valley–type (MVT) Pb–Zn deposits, and hydrocarbons is not well constrained. This is despite the fact that hydrocarbons generally occur in MVT deposits; the ores are emplaced in the same temperature range as hydrocarbon maturation and migration, and the deposits commonly occur in proximity to metal-rich black shales. Better understanding should lead to better exploration models for both hydrocarbons and MVT deposits. This connection is better understood with the help of Pb isotope patterns. Sphalerite Pb isotope compositions from the northern Arkansas and Tri-State mining districts and Woodford–Chattanooga and Fayetteville Shales were determined to assess the potential of shales as source rocks for the ore metals. The ores in both districts have a broad range of Pb isotope ratios and define linear trends, suggesting mixing of Pb from two distinct end members. Current results and previous depositional environment studies indicate the following: (1) shales deposited mainly under nonsulfidic anoxic conditions represent the less radiogenic end member, or (2) shales are the only source of ore metals. Given the array of organic molecules, each with their own thermochemical range, and the ways metals can be associated with them, the release of metals may cover varying ranges. Thus, the compositions of the released fluids would change through time and not have a single static composition, closely approximating the isotopic composition of the released metals at various times. Mineralization derived from a dynamically evolving fluid may show apparent end members, without the need to call on mixing of fluids from separate sources.
Mississippi Valley–type (MVT) deposits are important economic sources of both Pb and Zn, accounting for 38% of the global tonnage of all sediment-hosted Pb and Zn deposits (Leach et al., 2010). The most important minerals are sphalerite, galena, pyrite, marcasite, dolomite, and calcite; barite is commonly minor or absent, and fluorite is rare (Leach and Sangster, 1993; Leach et al., 2005). The MVT deposits are found worldwide, but many of the largest and initially studied deposits occur within the Mississippi River drainage of North America (Leach et al., 2010). Many of the MVT deposits formed during the Devonian to Permian, corresponding with intense tectonic events during assimilation of Pangea: the Acadian orogeny (Late Devonian–Early Mississippian), responsible for early mineralization in the Appalachian Mountain region, and the Alleghenian–Ouachita orogeny (Pennsylvanian–Permian) for mineralization in the Appalachian and Midcontinent regions (Leach et al., 2010; Gregg and Shelton, 2012).
The MVT ores are generally found in close proximity to hydrocarbon-bearing units and are often marked by the presence of hydrocarbon fluid inclusions and interstitial hydrocarbons, including heavy oil and bitumen (Sverjensky, 1984; Gregg, 2004). The MVT deposits are part of the development of sedimentary basins and are not clearly related to a crustal (igneous) source for the metals (Macqueen, 1986). They commonly occur in small to large deposits in carbonates around the edges of basins, often located where organic-rich shales onlap platform carbonate margins on the flanks of sedimentary basins (Figure 1; Leach et al., 2010; Schutter, 2015). Black shales associated with these sedimentary basins in the central United States include Devonian and Carboniferous shales: the Marcellus Shale (Devonian) and the Chattanooga Shale (Devonian–Mississippian) in the Appalachian basin, the Woodford Shale (Devonian–Mississippian) and the Fayetteville Shale (Mississippian) in the Arkoma basin, the New Albany Shale (Devonian–Mississippian) in the Illinois basin, the Antrim Shale (Devonian–Mississippian) in the Michigan basin, and the Pennsylvanian shales in the Forest City basin and the Midcontinent region (Oklahoma, Kansas, Missouri, Nebraska, Iowa, Illinois, and Indiana [Figure 1]). However, MVT Pb–Zn mineralization is not restricted to the margins of foreland basins. The Upper Mississippi Valley district (Figure 1) is far removed from any marginal fold belt.
Figure 1. Map of the United States Midcontinent region showing the location of sedimentary basins, shale plays, and major Mississippi Valley–type (MVT) mining districts and deposits (1 through 12): (1) northern Arkansas, (2) Tri-State, (3) Burkesville deposit, (4) Old Lead Belt, (5) Southeast Missouri Barite, (6) Central Missouri Barite, (7) Viburnum Trend, (8) Upper Mississippi Valley, (9) Illinois–Kentucky Fluorspar, (10) central Tennessee, (11) central Kentucky, and (12) eastern Tennessee. The represented shale plays are Antrim (Michigan Basin), New Albany (Illinois Basin), Excello–Mulky (Forest City Basin and Cherokee Platform), Woodford (Anadarko Basin), Barnett (Fort Worth Basin), Haynesville–Bossier (Texas–Louisiana–Mississippi [TX-LA-MS] Salt Basin), Fayetteville (Arkoma Basin), and Chattanooga (Black Warrior Basin and Appalachian Basin). Modified from Garven et al. (1993). Inset: map of the Arkoma foreland basin and the corresponding forebulge, known as the Ozark Dome, showing the location of the analyzed ore samples (Tri-State and northern Arkansas mining districts), basement (Spavinaw granite), and shale samples (Fayetteville Shale and Woodford–Chattanooga Shale). Modified from Bradley and Leach (2003). For drawing the basemap: https://www.eia.gov/maps/maps.htm; https://water.usgs.gov/GIS/metadata/usgswrd/XML/physio.xml. Miss. = Mississippian; Penn. = Pennsylvanian.
The major hydrocarbon resources may be concentrated elsewhere in the host basins from the sulfide ores (Ridley, 2013). However, as stated above, deposits exist where the MVT ores are directly associated with hydrocarbons (Sverjensky, 1984; Gregg, 2004). This suggests that the hydrothermal fluids and hydrocarbons were present together at the same time, implying that the MVT mineralization is associated with the generation and migration of hydrocarbons (Gregg and Shelton, 2012; Schutter, 2015). In some cases, the potential source rocks and their contained organics are thermally immature at the site of mineralization (Hatch et al., 1986; Macqueen, 1986). The presence of mature hydrocarbons within the fluid inclusions, despite the fact that the host rock itself is not thermally mature, implies that the hydrocarbons and ore-producing fluids migrated from deeper in the basin and were not matured locally (Schutter, 2015). Moreover, fluid inclusion studies in MVT ores suggest that the temperatures of the mineralizing fluid most commonly range between 75°C and 150°C (Leach et al., 2005). These values correspond with the oil generation window (between 50°C and 150°C [Tissot and Welte, 1984; Mastalerz et al., 2013; Schutter, 2015] and between 75°C and 175°C [Selley and Sonnenberg, 2014]). Many authors have commented on the coexistence and potential relationship between hydrocarbons and MVT ores. Veatch (1899) described sphalerite and galena found in conjunction with gas and oil at the crest of anticlinal salt mounds in Louisiana. Siebenthal (1915), in his review of the ore deposits around Joplin, Missouri, described heavy oil residue in contact with the base of the Pennsylvanian shales and in close proximity to the ore deposits. He interpreted this residue as evidence for ascending mineralizing fluids migrating with the bitumen. Dozy (1970) noted that the brines-connate or formation waters from adjacent basins, which might have supplied the metallic elements to the MVT ores in the midcontinent area of the United States, had also been the medium with which oil had been closely associated from its genesis through migration to accumulation. Hydrocarbons have been found in conjunction with Pb–Zn deposits in Sweden (Rickard et al. 1975), England (Parnell and Swainbank, 1990), Canada, and Australia. Oliver (1986) hypothesized that fluids mobilized from orogenies play a key factor in the migration of both hydrocarbons and Pb–Zn metals. Eisenlohr et al. (1994) linked the maturation and migration of hydrocarbons to the migration of mineralizing fluids responsible for the MVT ores in the Canning Basin of Australia. Understanding the link between the MVT ores and hydrocarbons can help clarify both the hydrocarbon systems and the occurrence of economic Pb–Zn deposits.
The main objectives of this study are the following: (1) to elucidate what factor, if any, the organic-rich Woodford–Chattanooga and Fayetteville Shales and the maturation of hydrocarbons play in the formation of the MVT deposits in the northern Arkansas and the Tri-State mining districts, and (2) to identify metal-contributing source rocks in the MVT sulfide ores from the southern Ozark region. To understand the linkage between MVT deposits and hydrocarbons, two lines of evidence have been considered. One is the distribution of mineralization in relation to the potential source rocks, and the second is the isotopic evidence. Although not ideal, to date, Pb isotopes represent the best available tracer of metal sources in hydrothermal systems. Their utility for constraining the source of associated metals Au, Ag, Cu, Zn, and other metals is limited by the assumption that Pb was derived from the same source, transported, and deposited from the same hydrothermal fluid (Tosdal et al., 1999). This assumption is mostly true, considering the similar geochemical behavior of Pb, Cu, and Zn in hydrothermal fluids (Henley et al., 1984) and the occurrence of Pb in the same paragenetic stages as Zn, Cu, Ag, and Cd (Bourcier and Barnes, 1987; Wood et al., 1987). However, for Au-only or Pb-poor systems, this assumption may not be valid because Pb may have been derived from rocks along the fluid channels (Powell et al., 1991). Nonetheless, the widespread occurrence of Pb in ore minerals deposited at all temperatures coupled with the lack of useful isotopic variations in other metals of economic interest makes Pb isotopes to serve as a proxy for the source(s) of associated metals (Tosdal et al., 1999). Given that Pb isotopes are not measurably fractionated by redox reactions in solution or by fluid–mineral interactions, the isotopic composition of Pb in ores is similar to that of Pb in its source (Tosdal et al., 1999). Here, the Pb isotope ratios of organic rich shales and granitic basement rocks are compared with the Pb isotope ratios of sulfide ores from the northern Arkansas and the Tri-State mining districts. To characterize the organic richness of the Woodford–Chattanooga Shale and Fayetteville Shale, the shale samples are subjected to an additional type of geochemical analysis, the total organic carbon (TOC) content.
The northern Arkansas and the Tri-State mining districts are situated on the southern edge of the Ozark Plateau and north of the Arkoma Basin (Figure 1). The Arkoma Basin is a deep Pennsylvanian foreland basin that was formed in response to crustal loading during the Ouachita orogeny (Viele and Thomas, 1989). The basin contains Upper Cambrian–Lower Pennsylvanian layers deposited on the shelf of a passive margin, overlain by Lower–Middle Pennsylvanian foreland basin deposits (Houseknecht et al., 2014). The Ozark Dome is a Precambrian (1.4 Ga) igneous-cored uplift, with the core being exposed in the St. Francois Mountains of southeastern Missouri. Paleozoic sedimentary rocks unconformably overlie the older igneous rocks and dip away in all directions from this igneous core. The Ouachita orogeny ended in the Late Pennsylvanian–early Permian (Arbenz, 1989), corresponding with the emplacement of the MVT ores in both districts (Pan et al., 1990; Brannon et al., 1996). The conventional MVT deposit model holds that the ores formed from very large hydrothermal systems in which the fluid drive was caused by deformation of foredeeps and uplift of foreland thrust belts during collisional tectonics (Leach and Rowan, 1986). The fluid flow was orogenically and topographically driven. Ground water, recharged in the uplifted orogenic margin of the foredeep, migrates through deep parts of the basin, acquires heat and dissolved components, and discharges along the basin’s cratonic flank (Garven, 1985; Bethke, 1986; Bethke and Marshak, 1990). The mobilized metals are carried by very dense Na–Ca–chloride brines, also called “oil-field” brines because they are saline waters typically found with oil in many basins (Ridley, 2013; Schutter, 2015). Much of the ore mineralization occurs as replacement of carbonate rocks and, to a lower degree, as open-space fill (Leach et al., 2010). The most important ore controls are faults and fractures, dissolution collapse breccias, and lithological transitions (Leach et al., 2005, 2010).
The Woodford–Chattanooga and the Fayetteville Shales
The Woodford–Chattanooga Shale is an Upper Devonian to Lower Mississippian shale that covers a wide swath of the North American continent from Oklahoma to Tennessee. It has long been regarded as an excellent hydrocarbon source rock in the Arkoma basin and other basins around Oklahoma (Byrnes and Lawyer, 1999) and a producer of both conventional and unconventional resources (Comer, 2005; Houseknecht et al., 2014; Infante-Paez et al., 2017). This formation generally thickens from north to south, ranging from less than 15 m (<50 ft) to greater than 91 m (>300 ft) thick in Oklahoma, and between approximately 15 m (∼50 ft) and 46 m (∼150 ft) thick in Arkansas (Comer, 1992; Li et al., 2010). It consists of gray to black, fissile, locally concretionary shale, and interbedded chert (Paxton et al., 2006). The shale corresponds with a major transgression across much of the area (Byrnes and Lawyer, 1999) and can be subdivided into three informal members (lower, middle, and upper), each differing in their depositional environments and lithological properties. The lower and middle members are progressively transgressive as the sea level rose; the upper member of the formation represents the maximum flooding interval when the shoreline began to prograde basinward (Slatt, 2013). The shale is very high in organic content, generally ranging between 1% and 10% TOC and locally exceeding 20% TOC (Carr, 1987; Comer and Hinch, 1987). Like many other organic-rich black shales, it is highly radioactive, to the extent of being considered a potential source of uranium (Glover, 1959; Swanson and Landis, 1962).
The Upper Mississippian Fayetteville Shale is a known source rock and producer of unconventional hydrocarbons within the Arkoma Basin (Ratchford and Bridges, 2006; Houseknecht et al., 2014). This shale is a basinal deposit produced by a northward expansion of the basin (Frezon and Glick, 1959). It has been interpreted as a transgressive facies deposited on the relict shelf, culminating in a maximum flooding surface in the upper Fayetteville (Handford and Manger, 1993; Houseknecht et al., 2014). The top of the Fayetteville Shale is a highstand deposit marking re-establishment of a carbonate ramp along the southern margin of the Ozark uplift (Houseknecht et al., 2014). The upper Fayetteville and lower Fayetteville are separated by the Wedington Sandstone Member in western Arkansas (Frezon and Glick, 1959). The upper Fayetteville is a black to gray shale with low organic content; the lower Fayetteville has a much higher organic content compared with the upper Fayetteville (Ratchford and Bridges, 2006). Present-day TOC generally ranges between 1% and 8% (Ratchford et al., 2006). The formation is thermally mature and produces dry gas across much of the Arkansas part of the northern Arkoma Basin (Ratchford and Bridges, 2006).
The Tri-State and the Northern Arkansas Mississippi Valley–Type Mining Districts
The Tri-State mining district is located in northeast Oklahoma, southeast Kansas, and southwest Missouri (Figure 1). The mineralized region covers a large area, approximately 200 km (∼125 mi) by 80 km (∼50 mi), and contains many scattered deposits of Pb and Zn ores (Brockie et al., 1968). The main ore minerals are sphalerite (ZnS) and galena (PbS), with sphalerite being more common than the galena by a ratio of approximately 5:1 (Brockie et al., 1968). The ore is in the Mississippian Boone Formation (Figure 2), a light-gray, crinoidal, finely crystalline limestone with variable proportions (20%–60%) of bedded and nodular chert (McKnight, 1935; Frezon and Glick, 1959; Snyder and Gerdemann, 1968). Every stratigraphic horizon of the Boone Formation has been mineralized to some extent. Shattered zones of rock are the most abundant host areas; local, brittle fractures create areas of excellent permeability and porosity, where mineralizing fluids later focused and deposited the ores (McKnight and Fischer, 1970). Normal faults are present but do not necessarily control ore locales, although fractured zones bordering the faults are important hosts for the ores. Many structural features are related to long-lived, deep-seated faults in the Precambrian basement rock that may have allowed deeper brines to rise into shallower carbonate strata (Brockie et al., 1968; Long et al., 1986). Based on Th–Pb dating of ore-stage calcite, main-stage mineralization in the Tri-State district has been constrained to the late Permian–Early Triassic (251 ± 11 Ma) (Brannon et al., 1996). Greater than 7,283,000 t of Zn ore (8,028,133 tons) and 176,600 t of Pb ore (194,668 tons) have been recovered from the area (Brockie et al., 1968).
Figure 2. Generalized stratigraphic sections of the units in the Ozark region. Stratigraphic locations of the ore deposits in Arkansas and the Tri-State region are indicated by black bars. Zigzag lines depict unconformities. Different colors have been used to depict various lithologies. Modified from Liner et al. (2013) and Wenz et al. (2012). CB = chert breccia; CH = chert; CT = chert tripolite; DOL = dolomite; FM = Formation; LS = limestone; SH = shale; SS = sandstone.
The northern Arkansas mining district is located in northcentral Arkansas (Figure 1). Mineralization developed along faults and occurs in solution collapse breccias and as replacement zones bordering faults and breccia zones (Leach et al., 1975; Plumlee et al., 1994). The sulfide mineralogy is mostly sphalerite (ZnS) with lesser amounts of galena (PbS), chalcopyrite (CuFeS2), pyrite (FeS2), marcasite (FeS2), enargite (Cu3AsS4), and greenockite (CdS) (McKnight, 1935). Most of the ores occur in the Ordovician Everton and the Mississippian Boone Formations (Figure 2). The Everton Formation is a blue-gray limestone interbedded with sandy limestone, sandstone, and sandy dolomite (McKnight, 1935; Snyder and Gerdemann, 1968). Similar to the Tri-State district, the rocks of the northern Arkansas district are not highly deformed or structurally complex, with only minor local folding and a low regional dip to the south (McKnight, 1935). Paleomagnetic studies (Pan et al., 1990) provide good constraints on the age of mineralization (Permian; 265 ± 20 Ma) in the northern Arkansas district. Between 1907 and 1930 production figures for Zn and Pb were 29,000 and 1,920 t (31,967 and 2116 tons), respectively (McKnight, 1935).
SAMPLES AND METHODS
Twenty-one sphalerite samples (11 from the northern Arkansas mining district and 10 from the Tri-State mining district), 13 Woodford–Chattanooga Shale samples, 10 Fayetteville Shale samples, and 2 granitic basement samples were analyzed for Pb isotope compositions. To characterize the organic richness of the shales, 12 Woodford–Chattanooga Shale samples and 7 Fayetteville Shale samples were analyzed for TOC content. The location of the samples is shown in Figure 1. Sphalerite ore samples were collected in the field from chat piles in the Tri-State mining district near Picher, Oklahoma (Figure 3A). Three other samples from the Tri-State mining district and the sphalerites from the northern Arkansas district (Figure 3B) were donated by D. L. Leach and M. S. Appold. The shale samples (Figure 3C, D) were collected in the field from outcrops in Arkansas, Oklahoma, and Missouri. Care was taken at the outcrop to dig back and collect as fresh and unaltered specimens as possible. The granitic samples were collected from an outcrop in Spavinaw, Oklahoma. The chemical processing of the rock and mineral samples was done in the modular Class 100 Radiogenic Isotope Laboratory at the University of Arkansas. We determined Pb isotope ratios using a Nu Instruments inductively coupled plasma mass spectrometer at the University of Arkansas. Because the concentration of U and Th, with respect to Pb, in ore minerals is intrinsically low (reported concentrations in sphalerite can be found in Gulson, 1984 and Goldhaber et al., 1995), time-integrated growth in the Pb isotope composition is minimal to negligible for minerals formed in the Phanerozoic (Tosdal et al., 1999). Therefore, the measured Pb isotope ratios of the sphalerites do not need to be age corrected for U and Th decay. The measured isotopic ratios approximate the composition of the mineral and the hydrothermal fluid at the time of crystallization if the system remained closed (Tosdal et al., 1999). In contrast, the measured Pb isotope ratios of the analyzed rocks need to be age corrected for U and Th decay to 250 Ma, the approximate age of the ores (Pan et al., 1990; Brannon et al., 1996). Age correction is necessary because U and Th are not excluded from the structure of common minerals found in many rocks, and therefore, time for radiogenic growth of Pb (time-integrated growth) exists. For calculating the initial Pb isotope ratios, we use U and Th concentration data reported by Paxton et al. (2008) for the lower Fayetteville Shale (6.7 ppm U and 14.1 ppm Th), the upper Fayetteville Shale (9.3 ppm U and 9 ppm Th), and the Chattanooga Shale (19.1 ppm U and 12.8 ppm Th) and Pb concentration data (20 ppm) reported by Condie (1993). The U and Th concentrations reported by Paxton et al. (2008) are quite close to the average U (2.7 ppm) and Th (12.3 ppm) concentrations reported by Condie (1993) for the North American shale composite. Total organic carbon analysis was performed on select shales using a Carlo Erba NC 2500 elemental analyzer (EA) at the University of Arkansas Stable Isotope Laboratory. A detailed description of the chemical processing of the samples, the Pb isotope analysis, and the TOC analysis is provided in Appendix 1 and Appendix 2.
Figure 3. Images showing ore samples from the Tri-State and the northern Arkansas Mississippi Valley–type mining districts (A and B), and the Fayetteville and Woodford–Chattanooga Shales (C and D).
Lead Isotope Background
Ordinary Pb (Russell et al., 1954) occurs in minerals whose U/Pb and Th/Pb ratios are so low that its isotopic composition does not change appreciably with time (Faure, 1986). The designation “J-type anomalous Pb” was proposed by Houtermans (1953) for Pb in galena from Joplin, Missouri, in the Tri-State mining district. Cannon et al. (1961) suggested two different processes that can generate this anomalous Pb: (1) evolution of Pb by accumulating increments of radiogenic Pb during a protracted span of geologic time, or (2) geochemical mixing of ordinary Pb with radiogenic Pb that formed in closed source regions having differing values of 238U/204Pb and 232Th/204Pb. The Pb isotope ratios of this anomalous Pb do not fit the single-stage model of Pb evolution (Holmes–Houtermans model, developed in 1946) and require interpretations by two-stage (Stacey and Kramers, 1975) or three-stage models (Goldhaber et al., 1995). The two-stage growth curve of Stacey and Kramers (1975) approximates the orogene growth curve of Zartman and Doe (1981). The single-stage model of Pb evolution assumes that radiogenic Pb is produced by decay of U and Th in the source regions (having different values of 238U/204Pb and 232Th/204Pb) and that the resulting Pb (primeval plus radiogenic) is then separated from its parents and incorporated into ore deposits as galena (Faure, 1986).
Typically, Pb isotope data are presented on two covariation diagrams: the thorogenic diagram (208Pb/204Pb vs. 206Pb/204Pb) and the uranogenic diagram (207Pb/204Pb vs. 206Pb/204Pb). The thorogenic diagram represents the radiogenic daughter of Th versus the radiogenic daughter of the most abundant U isotope; the uranogenic diagram represents the least abundant isotope of U versus the most abundant one (Tosdal et al., 1999). While in contact with aqueous fluids, Th is insoluble and one of the more inert elements (Taylor and McLennan, 1985). In contrast to Th, U has a second oxidation state; under oxidizing conditions, it forms uranyl ions (+6) that are highly soluble in aqueous fluids (Tosdal et al., 1999). Under these conditions, U may be significantly fractionated from Th. At low temperatures, Pb is easily complexed with organic matter and is generally insoluble (however, in hydrothermal environments, Pb is soluble).
Figure 4. Plots showing the 208Pb/204Pb versus 206Pb/204Pb (A) and 207Pb/204Pb versus 206Pb/204Pb (B) of sphalerite, basement, and shale samples analyzed in this study. Also shown in (B) are the slopes and coefficient of determination (R2) values of the linear trends defined by the northern Arkansas sphalerite samples. The Pb isotope ratios of the Spavinaw granite, and the Woodford–Chattanooga and Fayetteville Shales have been age corrected to 250 Ma. The average orogene (OR) growth curve is from Zartman and Doe (1981).
According to Zartman and Doe (1981), three idealized crustal reservoirs of U–Th–Pb exist: the mantle, lower crust, and upper crust. These reservoirs mix in the orogene, where crustal deformation, magmatism, sedimentation, and metamorphism occur (Tosdal et al., 1999). Model curves for each of these major sources are derived, and many Pb isotope studies of ore deposits discuss Pb isotope data with respect to these model reservoirs (Doe et al., 1979; Stacey and Hedlund, 1983; Puig, 1988; Beaudoin et al., 1992).
RESULTS AND DISCUSSION
The measured Pb isotope ratios of ores from the northern Arkansas and the Tri-State districts are presented in Tables 1 and 2 and plotted in Figure 4 alongside the orogene growth curve outlined in Zartman and Doe (1981). The orogene growth curve represents the Pb isotope composition of orogenic regions through time. On conventional Pb isotope diagrams, the ore samples from both districts plot beyond the 0 Ga Pb isotope values of the orogene growth curve (Figure 4). These results confirm previous studies (Cannon et al., 1961; Heyl et al., 1966; Deloule et al., 1986; Crocetti et al., 1988; Kesler et al., 1994a, b; Goldhaber et al., 1995; Potra and Moyers, 2017; Potra et al., 2018), showing that Pb in nearly all post-Precambrian MVT ores is enriched in the radiogenic isotopes (having 206Pb/204Pb ratios of 20 or greater) compared with ordinary (common) Pb.
As a group, the Tri-State ores are more enriched in the radiogenic Pb products than the northern Arkansas district ores, although a considerable amount of overlap exists (Figure 4). The overlap suggests that the Pb and Zn from both districts may share some common source. The ores in both districts have a broad range of isotopic ratios and define linear trends. Heterogeneous Pb isotope compositions are expected in sedimentary rock–hosted deposits where fluids may have traveled along different aquifers, equilibrated with rocks of different chemical and isotopic compositions, and mixed at the site of ore deposition (Tosdal et al., 1999). Conventionally, these linear trends have been interpreted to indicate mixing of Pb from two distinct end members. One end member must be highly radiogenic, with Pb isotope ratios equal or higher than the highest measured value of the ores. The other end member must be less radiogenic, with Pb isotope ratios equal or lower than the lowest measured value. Compositions of Pb isotopes of the ores will be intermediate between the two sources. Mixing of the fluids at or near the site of ore precipitation is required to maintain any Pb isotope distinction and is evidence for multiple fluid input (Tosdal et al., 1999). This observation falls in line with the observed trends in many MVT deposits in which the ores show extreme ranges of Pb isotope compositions, suggesting mixing between multiple reservoirs near the site of ore deposition (Heyl et al., 1966; Kesler et al., 1994b; Goldhaber et al., 1995). The mixing of fluids causes rapid changes in fluid chemistry, triggering metal precipitation. The mixing of fluids also occurs in zones of high porosity and high permeability, like the breccias, which commonly host the ore deposits. The ore samples from the Tri-State district analyzed in this study define a broader range of Pb isotope ratios compared with published Pb isotope ratios of ores in the same district (Potra et al., 2018). The author concluded that the narrow range of Pb isotope ratios defined by ores in the Tri-State district analyzed in their study might indicate that the mineralization process caused, to some extent, homogenization and mixing of the Pb that contributed to generation of the ores. This can occur if mixing of fluids occurred far away from the site of ore deposition. The distance required for Pb homogenization to occur is not yet known. The same authors concluded, as an alternative explanation, that the narrow range might indicate a single, well-mixed, and homogenized Pb source. It is possible that the Tri-State ores from the study (Potra et al., 2018) came from a more localized area within the Tri-State mining district compared with the samples analyzed in this study.
On both Pb isotope plots, the ore samples plot beyond the 0 Ga value of the orogene growth curve. However, they plot to the right of the curve on the thorogenic diagram (Figure 4A) and to the left of the curve on the uranogenic diagram (Figure 4B). This implies that, relative to the orogene growth curve, the ore samples show lower 208Pb/204Pb and higher 207Pb/204Pb for a given 206Pb/204Pb. This pattern is also noticed in other ore districts from the Ozark region (southeast Missouri and central Missouri districts; Figure 1), characterized by lower 208Pb/204Pb and higher 207Pb/204Pb values for a given 206Pb/204Pb compared with the Upper Mississippi Valley, Illinois–Kentucky, central and southern Appalachian, east and Middle Tennessee, and central Kentucky districts (Heyl et al., 1966; Deloule et al., 1986; Crocetti et al., 1988; Kesler et al., 1994a, b; Goldhaber et al., 1995; Potra and Moyers, 2017; Potra et al., 2018). Elevated 207Pb/204Pb values may be indicative of regions of the crust where radiogenic Pb evolved in Precambrian rocks because of the relative abundance of 235U in the Precambrian. Because of the short half-life of 235U (0.7 Ga) and its fast decay rate to 207Pb, the present abundance of 235U is only 0.7%. Lower 207Pb/204Pb values suggest derivation of that Pb at recent times from a mantle or an oceanic environment and then incorporation into a rock or mineral (Tosdal et al., 1999). However, high-grade metamorphism can also cause U loss in a rock, leading to low 207Pb/204Pb values.
The slope of the linear trend defined by the Pb isotope ratios of the northern Arkansas ores is 0.0747 (Figure 4B), which corresponds with an age of approximately 1.05 Ga. This implies input of Pb from a Precambrian source. Therefore, an alternative for the linear array is that the basement rocks supplied a percentage of the ore Pb. In this case, the observed linear array in the northern Arkansas ores represents a “pseudo-isochron” (Jones et al., 2017) inherited, but not entirely, from the local basement. The Spavinaw granite (1.35–1.40 Ga), which represents the middle Proterozoic basement (Bickford et al., 1981; Kisvarsanyi, 1990) underlying parts of Kansas, Oklahoma, Missouri, and Arkansas, is part of the southern granite–rhyolite province. Notably, the southeast Missouri district, located east-northeast of the study area, is underlain by the basement rocks of the St. Francois Mountains (1.42–1.48 Ga), which are part of the eastern granite–rhyolite province (Bickford et al., 1981; Goldhaber et al., 1995). These two provinces cover a large part of the midcontinent United States. The measured Pb isotope ratios of the granites were age corrected for U and Th decay to 250 Ma, the approximate age of the ores. The age-corrected Pb isotope ratios of the Spavinaw granite plot close to the 0 Ga value of the orogene growth curve (Figure 4), corresponding with a Pb isotope evolution curve for 238U/204Pb (or μ) of approximately 9.74 and for 232Th/228U (or κ) of approximately 3.8 (Zartman and Doe, 1981). The age-corrected Pb isotope ratios of the granite samples analyzed in this study plot close to the least-radiogenic end of the ores. However, relative to the orogene growth curve, the granite samples do not show the lower 208Pb/204Pb and higher 207Pb/204Pb for a given 206Pb/204Pb signatures characteristic to the ores, suggesting that the currently analyzed granites did not contribute Pb to the ores in the northern Arkansas and the Tri-State mining districts. Considering the low number of granites analyzed in this study, it is premature to preclude the involvement of basement rocks in the generation of ores in the southern Ozark region. In their ore genesis study of the southeast Missouri district, Goldhaber et al. (1995) indicated that the basement and basement-derived siliciclastics contributed varying proportions of Pb and other metals to MVT deposits in the Ozark region. The authors concluded that three stages of U enrichment were required in the source rocks to model the Pb isotope evolution of the ore Pb and that 238U/204Pb values of less than 20 to greater than 40 were necessary to explain the various ore Pb arrays.
The measured and age-corrected Pb isotope ratios of the Woodford–Chattanooga and Fayetteville Shales are presented in Tables 3–5. The measured Pb isotope ratios of the whole-rock shale samples were age corrected for U and Th decay to 250 Ma, the approximate age of the ores. Relative to the orogene growth curve of Zartman and Doe (1981), greater than half of the analyzed shales plot to the left of the curve on the thorogenic diagram (Figure 4A), suggesting that they were less likely to be involved in supplying the ore Pb. However, some of the analyzed shales plot close to the least-radiogenic end of ore samples on the same Pb isotope diagram (Figure 4A). Similar to the ore samples, the analyzed shales plot above and to the left of the orogene growth curve on the uranogenic diagram (Figure 4B), suggesting greater time-integrated U/Pb values compared with the Spavinaw granites. Overall, the age-corrected 207Pb/204Pb and 208Pb/204Pb ratios of the Woodford–Chattanooga and Fayetteville Shales are lower compared with the Pb isotope ratios of the southern Ozark ores. On the uranogenic diagram, the majority of the shale samples plot close to the least-radiogenic end of the ores (Figure 4B). These results imply that only some of the analyzed shales may represent the less-radiogenic end member that contributed Pb to the ores.
Model of Mineralization
An alternate model is that organic-rich shales are the only source of metals for MVT ores. Many organic-rich shales have a very high U and organic matter content (uraniferous oil shales), being considered a potential source of both oil and U (Swanson, 1960). They would, in turn, provide a source for radiogenic Pb to be mobilized into MVT deposits. Uranium content varies by depositional environment. Humic-type land organics have a higher affinity for U than the sapropelic-type marine organics do (Swanson, 1960; Coveney and Glascock, 1989). On a first look, this suggests that a black shale that is deposited closer to shore might have a higher U content than one farther from shore, but another possible variable exists; humic organics are more resistant to oxidation and degradation, so that as deposition slows and degradation progresses, the resulting organic sediments reaching the bottom are proportionally enriched in terrestrial material land more likely to pick up U (Schutter, 2016). A wide range of organic molecules in organic-rich shales (aliphatic, aromatic, and NSO compounds) exists. The black shales deposited in a nearshore or estuarine environment are enriched in land-derived humic organics constructed of aromatic (unsaturated-chain) building blocks. The black shales deposited offshore are enriched in marine organics, which contain more aliphatic (saturated-chain) building blocks. The metals can be associated with these very complex molecules in a variety of ways. Some metals are vital components of organic activity, such as Ca, Mg, and Fe. Some are micronutrients, present in very low amounts, but still necessary for specific reactions, such as Zn and Ni. Some follow other elements around and poison the main reactions, such as Pb and Ba following after Ca. Adsorbed or interstitially trapped metals also exist. Considering the range of organic molecules and the ways metals can be associated with them, a difference in the way the various black shales accumulate and release metals exists. As the organic matter in the black shales (with adsorbed metals) undergoes catagenesis during burial and heating (maturation of hydrocarbons), the organic molecules are progressively cracked, releasing the adsorbed metals. Some evidence that progressive catagenesis releases different cations in succession exists (Figure 5). For example, Zn seems to be available at fairly low temperatures; released Fe is available at higher temperatures (some of the sphalerite-replaced fossils have late ruby sphalerite filling voids within the pale-yellow sphalerite); Pb may be available at an even later and higher temperature (Figure 5). Thus, metal release and migration would be expected to occur in the same temperature range as hydrocarbon maturation and migration from the source rocks (Figure 6). Given the vast range of organic molecules, each with their own thermochemical ranges, the actual release of metals may cover varying ranges. Thus, the compositions of the released fluids would change through time, depending on the thermal history, closely approximating the isotopic composition of the released metals at various times. Therefore, fluids would evolve and not have a single static composition. This is unlike what happens during inorganic diagenesis, in which specific reactions occur at specific temperatures. Mineralization derived from a dynamically evolving fluid may show apparent end members, without the need to call on mixing of fluids from separate, distinct sources. If this is a valid model, it would be expected that mineralization, especially Pb and Zn mineralization, would reflect the distribution of metal-enriched organic-rich shales.
Figure 5. Conceptual diagram of increasing chemical activity with depth and temperature. The tracks for the various metals and hydrocarbons reflect the relative burial conditions when they are released from organic-rich black shales. The two tracks for Fe (short dash and long dash) indicate that with very high levels of available organic material (short dash), the available mobile iron may be depleted (Leventhal, 1998) and may not become available again until released diagenetically. Modified from Schutter (2015; figure 5).
The Pennsylvanian marine black shales of the Midcontinent (Oklahoma, Kansas, Missouri, Nebraska, Iowa, Illinois, and Indiana) are highly enriched in heavy metals, averaging 1300 ppm Zn, 655 ppm Mo, 1400 ppm V, 85 ppm U, and 55 ppm Cd (Coveney and Glascock, 1989). They are also enriched in Ag, Au, Ba, Cu, Hg, Ni, Re, Se, and platinum group elements (Heckel, 1977; Coveney, 1981, 2003; Algeo and Maynard, 2004). These shales have very high TOC contents, with variable composition of organic matter type (Algeo and Heckel, 2008). Many of the black shales are actually iron limited; far less pyrite is present than would be expected from the amount of TOC present (Raiswell and Berner, 1985; Dean and Arthur, 1989; Leventhal, 1998; Algeo and Maynard, 2004). This lack of iron is consistent with the iron-poor sphalerite distributed across the Midcontinent (Schutter, 2015). The Upper Pennsylvanian marine black shales, referred to as Heebner-type shales by Heckel (1977), are euxinic (free H2S present in the water column) shales that formed offshore, probably in deep water, euxinic, and starved-basin conditions (Coveney and Glascock, 1989). The Middle Pennsylvanian marine black shales, referred to as Mecca-type shales (Tourtelot, 1963), are anoxic (no free H2S present in the water column) shales that were deposited near an ancient shoreline characterized by abundant peat swamps and deltaic siliciclastic sediments (Coveney and Glascock, 1989).
Figure 6. Relationships between hydrocarbon maturity, temperature, and vitrinite reflectance. Adapted from Etminan et al. (1984; figure 2) and Mastarlarz et al., (2013; figure 17). Temperature range in Mississippi Valley–type deposits, indicated by sphalerite inclusion temperatures, is typically approximately 60°C to 160°C, occasionally as high as 250°C (as per Coveney and Goebel, 1981; Hatch et al., 1986; Spirakis, 1986; Anderson and Macqueen, 1990; Evans, 1993; Gregg and Shelton, 2012) and effectively the same range as the oil window. From Schutter (2015). Orange shading represents the determined maximum temperature and vitrinite reflectance (Ro) in inclusions within sphalerite samples. The shape represents the frequency of such determined values. HC = hydrocarbons.
Coveney (2003) and Schutter (2015) reported on these Pennsylvanian shales and the adjacent strata, suggesting the presence of visible sphalerite (ZnS) in more than 20% of the samples. The sphalerites occur mainly in the muddy carbonates adjacent to the shales, where they replaced shelly material. These sphalerites occur far away (hundreds of kilometers or miles) from the Ouachita fold belt and trough, but in close proximity (only centimeters or inches) to the metalliferous black shales. This was uniformly true from Oklahoma to Iowa as well as in Illinois. The sphalerites are very pale yellow and apparently formed at fairly low temperatures, possibly in the 60°C–70°C range (Figure 6), supported by the unaltered color of conodonts and spores in the samples as well as sparse published fluid inclusion data. This minimal level of heating may account for the low level of iron in the sphalerite, and is below the oil window of hydrocarbon maturation. According to Tourtelot (1963), the sphalerite in the Mecca-type black shales precipitated during the earliest stages of diagenesis when H2S, released by decaying organisms, reacted with Zn dissolved in pore fluids. In the Heebner-type shales, Coveney (1979) concluded that sphalerite precipitation occurred prior to compaction. The same author proposed that the connate brines leached Zn from specific Upper Pennsylvanian black shales (Heebner-type shales) and precipitated sphalerite in the subjacent Mississippian host rocks. More recent studies (Selleck, 2014) conducted on the Utica Shale (Upper Ordovician) and Marcellus Shale (Middle Devonian) also note the presence of early diagenetic sulfides, including sphalerite. Therefore, it appears that only some sphalerites were deposited early in the organic-rich shales, favoring a syngenetic mineralization, whereas others involved a hydrothermal mineralization. This mineralization is very difficult to explain by the conventional MVT model. Considering the total amount of Zn involved (covering hundreds of kilometers, or miles, in lateral distance and hundreds of meters, or feet, of vertical section), deriving it all from a distant deep basin becomes problematic. Additionally, considering that the Zn would need to travel hundreds of kilometers, or miles, through poorly porous rock and then become preferentially precipitated in muddy carbonates, a distant source becomes even more unlikely. The Zn-rich occurrence in southwest Iowa is in the Forest City Basin (Figure 1), a completely separate basin that does not represent a foredeep. Some of the highest Zn concentrations occur in Forest City Basin samples. Therefore, it is much more likely that most of the metals were sourced from the adjacent black shale rather than a presumed distant basin. The isotope data also tell a compatible story. As mentioned above, Pb in MVT ores is highly radiogenic (J-type Pb). If the occurrence of Pb is dependent on hydrothermal sources, then J-type Pb is a coincidence, depending on an undocumented U-rich source, generally attributed to the crystalline basement. If, however, the occurrence of radiogenic Pb is a function of a source rock enriched in U, then it becomes a “fingerprint” of the base metal and hydrocarbon system.
The TOC values for the Woodford–Chattanooga Shale samples range from 2.16% to 7.51% (Tables 6, 7), indicating their high potential as hydrocarbon source rocks. Values for the Fayetteville Shale samples range from 0.09% TOC in the uppermost section to 7.48% TOC in the lower section (Tables 6, 7), with the majority also having potential for great hydrocarbon source rocks. The TOC analysis performed on the shales indicates that most samples have a high TOC content, with the exception of the samples from the upper member of the Fayetteville Shale. The shale samples that are enriched in the uranogenic Pb also have the highest TOC content (Figure 7A). For the thorogenic Pb, however, this pattern is only noticed in the Fayetteville Shale, but not in the Chattanooga Shale (Figure 7B). The high uranogenic Pb can be explained by elevated U levels associated with increased preservation of organic matter in an anoxic depositional environment. Many authors have commented on the correlation between elevated U content and elevated TOC in shales. Uranium radioactivity content from a spectral gamma ray log is commonly used as a proxy for organic content of the rock and is also used to calculate TOC for evaluation purposes (Passey et al., 1990; Lüning and Kolonic, 2003; Gonzalez et al. 2013; Huang et al., 2015). The elevated U content is a result of reduced U being much less soluble than oxidized U. Because organic matter acts as a powerful reductant, anoxic environments elevate U concentrations in sediments by preventing its oxidation and preserving organic matter. The different behavior of Th compared with U is primarily because of the difference in sensitivities of these elements to oxygenation, which causes differences in their mobility in aqueous systems (Klinkhammer and Palmer, 1991). This separation can be further enlarged by differences in materials that allow fixation of U and not Th, and vice versa (Adams and Weaver, 1958), which potentially explains the trend noticed in the analyzed samples.
Biomarker studies have also been used to assess the paleodepositional environments of each member of the Woodford–Chattanooga Shale. Chlorobiaceae, or green sulfur bacteria, are indicative of euxinic (anoxic and sulfidic) depositional environments and a stratified water column. The bacteria require light and euxinic water conditions to live (Frigaard and Dahl, 2008). Evidence for these bacteria have been found in both the lower and middle members of the Woodford–Chattanooga Shale (Connock et al., 2018), but not throughout all members. No evidence suggesting that the upper members of the Woodford–Chattanooga Shale were deposited under euxinic conditions exists. The Fayetteville Shale was deposited under anoxic, but not euxinic, conditions (Ratchford and Li, 2008; Alase, 2012; Ceron and Slatt, 2012). This is significant because it implies that euxinic conditions could have persisted during the deposition of the basal Woodford–Chattanooga Shale. Exposure of the Zn chloride complex to a source of reduced sulfur (H2S) destabilizes the complex and triggers mineral precipitation.
Thus, it appears that the euxinic waters create the prime conditions for ore precipitation in the shales, thus reducing the availability of free Zn2+. However, under nonsulfidic anoxic conditions, the availability of free Zn2+ is higher compared with euxinic conditions. The anoxic conditions that prevailed during the deposition of several shale members of the Woodford–Chattanooga and Fayetteville Shales, coupled with geochemical data, suggest that mainly the shales deposited under nonsulfidic anoxic conditions pose as potential sources of ore metals. As the shales undergo burial and maturation, Zn and Pb are mobilized, with adsorbed metals being released from cracking organics. Precipitation of ores requires the presence of reduced sulfur, and this is when the factor of hydrocarbons have come into play. The organic matter may have acted as a reductant and reduced sulfur from sulfate ions, or H2S from the maturation of hydrocarbons may have come in contact with the metal-rich fluids. However, these results do not exclude the euxinic muds as possible ore metal sources. Historically, the metal enrichment in black shales has been considered to be in base metal sulfides, as it occurs today. Redistribution into ore bodies may have required mobilization of black shale sulfides and redeposition in ore bodies. The current Pb isotope results and previous depositional environment studies described above indicate that the MVT ores may not be tied to the shales in general, but instead are favored by a particular environment within the larger black shale body. Uranium content is dependent on organic matter type (Swanson, 1960), and so, shales vary laterally and vertically in their original U content (Schutter, 2016) and, therefore, in their potential to release radiogenic Pb. The Pb might have been mobilized from shale when adequate thermal conditions were reached.
Figure 7. Plot showing the total organic carbon (TOC) content versus measured 206Pb/204Pb (A) and 208Pb/204Pb (B) of the Fayetteville and Chattanooga Shales.
Although the foregoing arguments broadly establish that the Fayetteville and Woodford–Chattanooga black shales may be the source of the mineralization in the Tri-State and northern Arkansas mining districts, part of the story remains to be considered, especially for its broader implications for MVT deposits in general. It is highly possible that the organic fraction of the shales yields distinct Pb isotope signature and U–Th–Pb concentrations compared with the aluminosilicate part of the shale. In this study, however, Pb isotope analyses were carried out only on the bulk shale samples. To better assess their involvement in the ore generation, a future work will include Pb isotope and U–Th–Pb concentration analyses of the following shale fractions: (1) insoluble organic matter (kerogen concentrates), (2) extractable organic matter (bitumen concentrates), (3) leachate part following extraction of organic matter (partial chemical dissolution), and (4) residue part following extraction of organic matter (total chemical dissolution).
This study potentially explains several issues related to MVT deposits.
- Why MVT mineralization is accompanied by hydrocarbons.
- Why MVT minerals occur in the same temperature range as the oil window of hydrocarbon generation (Figure 6.
- The widespread distribution of sphalerite (MVT-related) mineralization reflects the distribution of metal-rich, organic-rich shales, not deep basins.
- The metal and hydrocarbon release can be expressed as an evolving system and does not necessarily require the mixing of multiple fluids from different sources.
- The variable organic matter type within a shale (Schutter, 2016) may account for variable Pb isotope signatures in evolved fluids and cations.
- Provide a source highly enriched in U to be a source for radiogenic J-type Pb.
On conventional Pb isotope diagrams, the ore samples from the northern Arkansas and the Tri-State mining districts plot beyond the 0 Ga value of the orogene growth curve of Zartman and Doe (1981), confirming previous studies that show that Pb in nearly all super-Precambrian MVT ores is enriched in the radiogenic isotopes (having 206Pb/204Pb ratios of 20 or greater) compared with ordinary Pb. However, the ore samples plot to the right of the curve on the thorogenic diagram and to the left of the curve on the uranogenic diagram. This implies that, relative to the orogene growth curve, the ore samples have lower 208Pb/204Pb and higher 207Pb/204Pb for a given 206Pb/204Pb. They display a broad range of Pb isotope ratios and define linear trends, which suggest mixing of Pb from two distinct end members.
Relative to the orogene growth curve of Zartman and Doe (1981), greater than half of the analyzed shales plot to the left of the curve on the thorogenic diagram. These shales may have been deposited under euxinic conditions. However, some of the analyzed shales plot close to the least-radiogenic end of the ores and may have been deposited under nonsulfidic anoxic conditions. Similar to the ore samples, the analyzed shales plot above and to the left of the orogene growth curve on the uranogenic diagram. Overall, the age-corrected 207Pb/204Pb and 208Pb/204Pb ratios of the Woodford–Chattanooga and Fayetteville Shales are lower compared with the Pb isotope ratios of the southern Ozark ores. On the uranogenic diagram, the majority of the shale samples plot close to the least-radiogenic end of the ores. The current Pb isotope results and previous depositional environment studies indicate that shales deposited mainly under nonsulfidic anoxic conditions represent the less-radiogenic end member that contributed metals to the MVT ores. The more-radiogenic end member, though not fully supported by the current results (possibly because of the limited number of analyzed samples), may be represented by the basement rocks. However, an alternate model is that organic-rich shales are the only source of ore metals. Considering the range of organic molecules and the ways metals can be associated with them, a difference in the way the various black shales accumulate and release metals exists. As the organic matter in the black shales undergoes catagenesis during burial and heating (maturation of hydrocarbons), the organic molecules are progressively cracked, releasing the adsorbed metals. Given the vast range of organic molecules, each with their own thermochemical range, the actual release of metals may cover varying ranges. Thus, the compositions of the released fluids would change through time, depending on the thermal history, closely approximating the isotopic composition of the released metals at various times. Mineralization derived from a dynamically evolving fluid may show apparent end members without the need to call on mixing of fluids from separate, distinct sources.
Input of reduced sulfur is required for ore precipitation, and hydrocarbons may contribute here. The organic matter may have acted as a reductant and reduced sulfur from sulfate ions, or H2S from the maturation of hydrocarbons may have come in contact with the fluids. Given that these shales are great hydrocarbon source rocks and that the ores are commonly found in proximity to hydrocarbons, it is possible that the hydrocarbons and the metal-charged fluids migrated together and precipitated when conditions required it. An important implication is that if MVT deposits and hydrothermal maturation in marine source rocks are aspects of the same process, then they can inform and improve each other. Understanding MVT deposits can lead to enhanced knowledge about the presence and relevant characteristics of metal-bearing black shales (aka source rocks or unconventional resources). Conversely, understanding these black shales may lead to improved exploration for MVT deposits and their characteristics, particularly if they are temperature dependent. The MVT deposits, and even dispersed mineralization, can be built into conventional hydrocarbon basin modeling, improving both sides.
It is worth noting that many of the metals commonly occurring with hydrocarbons and in associated brines also occur in metal-rich, organic-rich shales. They may be mobilized by processes similar to those discussed here for Pb and Zn. Thus, understanding the mechanism can lead to modeling and predicting where they may be concentrated, possibly even leading to economic concentrations, or at least anticipating potential problems. Not all Pb–Zn deposits are MVT deposits; some are more conventional hydrothermal deposits, and some are ambiguous in origin. Clearly, much can be learned by resolving how the various mechanisms operate. Several types of other metal deposits that involve organic-rich black shales also exist. It is not clear if they are similar to MVT deposits or simply reflect hydrothermal systems interacting with organic-rich shales. Even if the latter is the case, study can enhance understanding of those shales and their structure.
APPENDIX 1: CHEMICAL PROCESSING OF THE SAMPLES
Chemical Processing of Sphalerite Samples
Fresh sphalerite crystals were handpicked from the ores and soaked for 30 min in nitric acid. The samples were rinsed with triple-distilled water and dried on a hot plate. Approximately 150 mg of each sample was weighed out, and 2 ml of 8 N HNO3 was added to each sample. The samples were then heated at 150°C and dried down on a hot plate within the laminar flow hood. Full digestion of undissolved samples was achieved by successive additions of 1 ml of 8 N HBr and 1 ml of 8 N HNO3. The samples were dried down at 150°C. Two milliliters of 1 N HBr was added to each sample and dried down. This step was repeated two more times. Five hundred microliters of 1 N HBr was added to each sample to redissolve them. Each sample was centrifuged in a HNO3-leached centrifuge tube for 10 min, rotated 180°, and centrifuged for an additional 10 min. The samples were transferred to a 3-ml cation exchange column. Following the method outlined by Manhes et al. (1978), we separated and purified Pb. The column had 0.1 ml of Dowex AG1-8X 200–400 mesh resin with polytetrafluoroethylene (PTFE) frits; the resin was precleaned by mixing it with 6 N HCl and rinsing with 0.5 N HNO3 and triple-distilled water. The sample was washed with three successive additions of 1 ml of 1 N HBr, eluted using 1 ml of 20% HNO3 into a 5-ml PTFE container, and dried in a laminar flow hood.
Chemical Processing of Shales
Each specimen was rinsed in deionized (DI) water and allowed to dry. The samples were wrapped in paper towels and aluminum foil, placed in plastic bags, and crushed with a hammer. Fresh, unaltered chips of rock were collected from the crushed rock samples. The selected chips were powdered using a SPEX SamplePrep ShatterBox. Between each sample, the alumina ceramic grinding container was cleaned with DI water, double-distilled water, and pure quartz sand to avoid cross-contamination. The quartz sand had been previously acid cleaned in HNO3, added to the container, and powdered to aid the cleaning process. Following this step, the vessel was self-contaminated by adding a few rock chips, powdering them, and then discarding the powdered sample. The remainder of the sample was powdered, and 250 mg of each rock sample was weighed out for isotopic analysis.
A high-purity, high-strength mixture of 5 ml of HF, 3 ml of HNO3, and 1 ml of HCl was added to each shale sample in accordance with the clay dissolution method outlined in the MARS laboratory microwave procedures. The shale samples were then heated in a MARS laboratory microwave. The temperature was ramped up to 200°C over a span of 15 min and then held at 200°C for 10 min. The shales were still not fully digested; therefore, they were heated again, this time ramping the temperature to 200°C over 15 min and holding it at 200°C for 45 min. The samples failed to fully digest, with the highly refractory minerals and elements that formed insoluble fluorides (e.g., Al, Ba, Ca, and Mg) remaining undissolved. However, this study does not focus on analyzing the aforementioned elements, and therefore, the process outlined above is reliable for the scope of the current research. Following the dissolution procedure, the shale samples were dried down and transferred to cation exchange columns. Following the method outlined by Pin et al. (2014), Pb was separated and purified.
Chemical Processing of Granites
Two granite samples were analyzed for their isotopic ratios. A total of 4 ml of 7 N HNO3 and 3 ml of HF were added to the granite and coprocessed blank sample and placed on a hot plate with the caps on the beakers until digested. Once digested, the caps were removed, and the samples were dried down. Successive additions of 0.5 ml 6 N HCl + 0.5 ml 7 N HNO3, 2 ml of HNO3, and 1 ml of HNO3 ensured full digestion of the igneous rock samples. Between each addition, the sample solutions were dried down. Two milliliters of 1 N HNO3 were added to the final dried samples, centrifuged for 15 min, rotated 180°, and centrifuged for an additional 15 min. The samples were transferred to cation exchange columns, and Pb was separated and purified following the method outlined by Pin et al. (2014).
APPENDIX 2: CHEMICAL ANALYSIS OF PROCESSED SAMPLES
Lead Isotope Analysis
The samples processed for Pb isotope analysis were diluted with 2% HNO3 containing 4 ppb Tl prior to analysis. The sample was introduced into the plasma by an uptake system with a rate of 40 μl/min. The aerosol from the nebulizer is injected into the center region of the plasma, desolvated, and ionized. The Pb–Tl mixtures were normalized using the Tl normalization technique for mass bias correction following the procedure of Kamenov et al. (2004). The data collected for each sample represented averages of 60 ratio measurements each.
Forty-seven analyses of the National Bureau of Standards 981 Pb standard yielded the following results: 208Pb/204Pb = 36.6772 (±0.0097 2σ), 207Pb/204Pb = 15.4847 (±0.0022 2σ), and 206Pb/204Pb = 16.9312 (±0.0027 2σ). All standard and sample Pb data were normalized to 205Tl/203Tl = 2.38750, and the raw data were corrected using the sample-standard bracketing method and accepted values for the Pb standard from Todt et al. (1996). Four duplicates (FS1L, FS6L, FS8UU, and FS10L) were prepared and analyzed to evaluate the accuracy and reproducibility of the measurements (Tables 3–5). Blanks were processed to determine the amount of Pb introduced in the samples during chemical processing. The total procedural blanks level for Pb was less than 10 pg.
Total Organic Carbon Analysis
The shales were dried and weighed into tin capsules containing approximately 100 μg of carbon. Following these steps, the samples were combusted at 100°C in a stream of helium to quantitatively produce CO2. The combustion gases were separated on a 4-M C/N column. The Carlo Erba NC 2500 EA was interfaced with a Delta Plus isotope ratio mass spectrometer (MS) via a ConFlo II interface. The MS simultaneously monitored masses 44, 45, and 46 during the analysis. A pure-gas CO2 reference pulse was admitted to the MS after the sample peak to generate the raw instrumental results. The raw results were normalized using standards to the Vienna Peedee belemnite scale.
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The authors thank Erik D. Pollock and Lindsey L. Conaway for their lab support during measurements by Nu Instruments inductively coupled plasma mass spectrometer and Carlo Erba NC 2500 elemental analyzer at the University of Arkansas. The authors also thank John Gist, Laynie Hardisty, and Thomas Liner for their field support, and D. L. Leach and M. S. Appold for donating the ore samples from the northern Arkansas and the Tri-State districts. We are grateful to the AAPG Bulletin reviewers for improving the manuscript through their insightful comments and critical reviews.