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ABSTRACT

The Mississippian limestone is a petroleum exploration target in northern Oklahoma, and diagenetic events are significant factors in controlling porosity. In this study, paleomagnetic data, supported by petrographic results, were used to determine the origin and timing of diagenetic events in five unoriented cores from northern Oklahoma. Petrographic analysis indicates a complex paragenetic sequence, which includes precipitation of sphalerite and baroque dolomite. Thermal demagnetization removes a low-temperature viscous remanent magnetization (VRM) and a chemical remanent magnetization (CRM) in magnetite. An attempt was made to orient the cores using the VRM but this resulted in a streaked distribution of declinations. The inclinations of the CRM in the specimens in the five cores are similar (mean = −2.6°) and the age of the CRM was determined by comparing the inclinations with the expected inclinations for the study area. This indicates remanence acquisition in the Permian (~310–290 Ma). This is consistent with dates for mineralization in the nearby Tri-State MVT deposit and for a hypothesized Permian hydrothermal alteration event in the study area. The age of the CRM and the presence of sphalerite and baroque dolomite suggest that the CRM was acquired via hydrothermal fluids in the Permian.

INTRODUCTION

Determining the origin and timing of diagenetic events is important in studies of basin evolution as well as for the development of hydrocarbon exploration strategies. The Mississippian limestones in northern Oklahoma and southern Kansas are a major hydrocarbon exploration play. The main reservoirs within this play include porous chert zones, spiculitic limestones, and cherty limestones (Montgomery et al., 1998; Franseen, 1999; Rogers, 2001; Mazzullo et al., 2009, 2011). Two models for origin of the porosity have been proposed. One model suggests that the porosity in the unit was created almost exclusively during early diagenesis (e.g., Rogers et al., 1995; Montgomery et al., 1998; Watney et al., 2001; Franseen, 2006; Mazzullo et al., 2009) and is related to dissolution by meteoric fluids at unconformities, which formed as a result of relative sea-level changes. A second model proposes that later hydrothermal events enhanced porosity and were a major driver of diagenesis (e.g., Young, 2010; Goldstein and King, 2014). The objective of this study is to constrain the timing and origin of diagenetic events in five cores in the Mississippian limestones in northern Oklahoma (Figure 1) using an integrated paleomagnetic and diagenetic approach. In particular, we will test if there is a chemical remanent magnetization (CRM) in the rocks that could be related to the hydrothermal or other events in terms of timing and petrographic evidence.

Figure 1.

Location map showing the cores sampled in this study (modified from Oklahoma Geological Survey map).

Figure 1.

Location map showing the cores sampled in this study (modified from Oklahoma Geological Survey map).

The absolute dating of diagenetic events can be difficult to determine. Paleomagnetic analysis is one approach that can provide dates by isolating CRMs that are acquired when authigenic magnetic minerals precipitate (e.g., Elmore et al., 2012). The pole position for the CRMs can be compared to the apparent polar wander path (APWP) to establish an absolute date. Various diagenetic events have been shown to cause precipitation of magnetic minerals, including burial processes such as the smectite to illite transformation (Katz et al., 2000; Tohver et al., 2008) and maturation of organic matter (Banerjee et al., 1997; Blumstein et al., 2004), as well as externally derived fluid flow events (e.g., Elmore et al., 1993; Dennie et al., 2012). Fracture contact tests (e.g., Dennie et al., 2012), petrographic, geochemical, burial history, and timing data can be used to test the origin of the CRMs (e.g., Elmore et al., 2012).

Many recent and ongoing studies are focusing on the diagenesis of the Mississippian Limestone (e.g., Montgomery et al., 1998; Watney et al., 2001; Mazzullo et al., 2009; Matson, 2013; Goldstein and King, 2014). As a result, the focus of this study will be on paleomagnetic analysis to date events. Petrographic and geochemical data will be utilized to provide information on the origin of the magnetizations. Establishing a paragenetic sequence in conjunction with geochemical and paleomagnetic results will provide further insight into the diagenetic mechanisms involved in this complex unit.

Paleomagnetic studies require oriented samples to determine a pole position and establish an absolute date for a CRM. If a core is not oriented, the viscous remanent magnetization (VRM) method can be used to orient the core although relatively few studies have used this approach (e.g., Van der Voo and Watts, 1978; Suk et al., 1993; Dennie et al., 2012). The cores that were available in this study are unoriented, and as a result, we will assess the VRM method to orient the cores.

GEOLOGIC SETTING

The study area is located in a starved basin, south of the Burlington shelf or ramp margin (Franseen, 1999; Watney et al., 2001), and south and west of the Nemaha uplift (Figure 1). In the study area, there are four epochs within the Mississippian during which deposition was occurring: Kinderhookian, Osagean, Meramecian, and Chesterian (Northcutt et al., 2001; Mazzullo et al., 2011).

Major and minor unconformities played a major role in the Mississippian limestone, influencing diagenesis and thickness of the individual units. Kinderhookian through Osagean unconformities are controlled primarily by eustatic sea-level changes during tectonic quiescence (Frezon and Jordan, 1979). A number of workers have recognized shallowing upward cycles, which are capped by unconformities (e.g., Watney et al., 2001; Mazzullo et al., 2009; Leblanc, 2014). Increased subsidence of the Anadarko and Ouachita basins to the south during the Meramecian suggests the possibility of some noneustatic control of deposition (Frezon and Jordon, 1979). During Chesterian time, the study area was influenced by regional tectonics from uplift of the Ozark dome in north–east Oklahoma (Frezon and Jordon, 1979).

During the Kinderhookian, limestone, shaly limestone, and locally glauconitic shaly siltstones were deposited (Northcutt et al., 2001; Mazzullo et al., 2011). Post-Kinderhookian erosion removed much of the section throughout Oklahoma, especially in central and southern parts of the state (Northcutt et al., 2001). During the Osagean, most of Oklahoma was under shallow, warm, well oxygenated marine waters, with limestone as the dominant deposit (Northcutt et al., 2001). Mazzullo et al. (2009) identifies these units as the Burlington–Keokuk Limestone and the Reed Springs Formation, topped by a prominent syn- or post-Osagean unconformity. Meramecian and Chesterian seas were a very similar environment to that of the Osagean, resulting in extensive limestone deposition, with episodic terrigenous input. Post-Mississippian uplift and erosion along the Nemaha uplift has removed much or all of the Meramecian and Chesterian strata (Northcutt et al., 2001).

The most prominent regional structures such as the Nemaha uplift, and the Wichita and Ouachita mountains, are the result of Pennsylvanian tectonic events. The uplift and flexure associated with these events are associated with the major Mississippian–Pennsylvanian unconformity (Frezon and Jordan, 1979). This unconformity has a discontinuous porous chert zone located just below, in the uppermost Mississippian section (Rogers, 2001). This interval is interpreted to have formed by subaerial exposure because of uplift or eustatic sea-level fall, followed by erosion and redeposition of the cherty limestones (Rogers, 2001). Rogers (2001) proposed that the cherty limestone was eroded, redeposited, and subsequently silica replaced calcite. This was followed by infiltration of meteoric waters during subaerial exposure, which dissolved any remaining calcite.

The likely source for the hydrocarbons in the Mississippian limestone was the Woodford Shale (Matson, 2013) although some intervals such as the Cowley facies may be a source rock (Watney, 2014). Oil started to be generated from the Woodford Shale in the Anadarko Basin to the south of the study area at about 335 Ma (Gaswirth and Higley, 2013).

Several diagenetic studies of the Mississippian limestone have been completed in northern Oklahoma and southern Kansas. Two models have been proposed for the origin of the porosity in the Mississippian limestone. The first model relates porosity to dissolution by meteoric fluids at unconformities (e.g., Rogers et al., 1995; Montgomery et al., 1998; Watney et al., 2001; Mazzullo et al., 2009). A second model suggests that late hydrothermal fluid flow enhanced porosity in the reservoirs (e.g., Young, 2010; Goldstein and King, 2014; Ramaker et al., 2014). Goldstein and King (2014) investigated Ordovician, Mississippian, and Pennsylvanian rocks, proposing that hydrothermal fluids enhanced reservoir porosity in the Mississippian limestone during the Permian and perhaps during Laramide deformation. A number of workers have also reported evidence for alteration by hydrothermal fluids near the study area (e.g., Banner et al., 1988a; Banner et al., 1988b; Coveney, 1992). Mississippi Valley-type deposits in the nearby Tri-State Mineral District provide evidence for hydrothermal fluids (Coveney, 1992). Radiometric dating of these deposits yields an age between late Permian and Early Triassic (Brannon et al., 1996), and paleomagnetic dating agrees with this timing (Pan et al., 1990).

The Mississippian limestone is a heavily faulted and fractured unit. Fractures identified by geophysical studies are interpreted to be the result of subaerial karsting and collapse, hydrothermal alteration and dissolution, and fault tectonism (Elebiju et al., 2011). Fractures are altered diagenetically and were likely filled with porous chert, forming a conduit for fluid flow (Elebiju et al., 2011).

METHODS

Slabs from five cores (A. F. Severin, Richard Chelf #1, Tubbs #3, Shell Sutton #1–16, and Shell Oldenberg #1–16) were described at the Oklahoma Petroleum Information Center (OPIC), Oklahoma Geologic Survey, Norman, Oklahoma, and samples were collected for thin sections and paleomagnetic analysis. The cores are Osagean except for the lower part of the Severin core, which is Kinderhookian in age. More than 100 polished thin sections were examined with a Zeiss Axio Imager.Z1 petrographic microscope, allowing for reflected light and transmitted light microscopy. A modified Dunham’s (1962) classification scheme was used for all facies descriptions and both depositional, and diagenetic features were noted in the thin sections. Thin sections and rock chips were examined using a scanning electron microscope (FEI Quanta 200) with an Oxford Instruments X-act Energy Dispersive Spectroscopy system at the University of Oklahoma Mewbourne College of Earth and Energy IC3 lab.

Cores for paleomagnetic analysis were collected on an average spacing of 0.3–1.2 m (1–4 ft). Additional samples were collected from such features such as mineralized fractures and altered intervals. One hundred twenty-six core plugs were collected from the A. F. Severin core, 89 plugs from the Richard Chelf #1 core, 45 plugs from the Tubbs #3 core, 47 plugs from the Shell Sutton #1–16 core, and 49 plugs from the Shell Oldenberg #1–16 core. Plugs were sampled at OPIC from available core butts using a water-cooled drill press with a nonmagnetic bit. These plugs were then cut to a standard 2.2 cm (0.9 in) length with an ASC Scientific dual blade saw, yielding two to four specimens. Specimens were then labeled as either coming from the innermost part of the core, or from the edges, to test for a drilling induced magnetization (DIRM).

Natural remanent magnetizations (NRMs) were measured using a 2G-Enterprises cryogenic magnetometer with DC SQUIDS within the magnetically shielded paleomagnetic laboratory at the University of Oklahoma. The specimens were subjected to stepwise thermal demagnetization using 24 steps from 40°C to 500°C using an ASC Scientific Thermal Specimen Demagnetizer. Some specimens were subjected to two low-temperature demagnetization (LTD) treatments prior to thermal demagnetization to test if the treatment was necessary to remove a component residing in multidomain magnetite (e.g., Dunlop and Argyle, 1991). The specimens were subjected to two LTD treatments in which they were immersed in liquid nitrogen and allowed to warm to room temperature in a zero magnetic field and then remeasured. The data were examined for changes in intensity and direction.

Because the cores were unoriented, an attempt was made to orient the core using the VRM (Van der Voo and Watts, 1978; Suk et al., 1993). The VRM is expected to align with the modern direction. The VRM was isolated for the specimens, and the declination was rotated to align with the known modern field. The characteristic remanent magnetization (ChRM) was rotated by the same amount.

Thermal demagnetization data were analyzed with Super-IAPD 2000 software and plotted on Zijderveld (1967) diagrams. Best-fit lines using principal component analysis (Kirschvink, 1980) with a minimum of four points were used to define the ChRM. All chosen components have a mean angular deviation (MAD) less than 15°, although most were less than 10°. Fisher (1953) and “inclination only” (e.g., McFadden and Reid, 1982) statistics were used to compute mean directions.

The magnetic mineralogy of 24 representative specimens was investigated using isothermal remanent magnetization (IRM) acquisition and thermal decay measurements. The samples were first subjected to an alternating field (AF) of 120 mT demagnetization, and then the specimen was measured. Next, the samples were subjected to an IRM with 25 steps between 10 and 2,500 mT using an ASC Scientific Impulse Magnetizer. The IRM acquisition data from three of the cores (A. F. Severin, Richard Chelf #1, and Tubbs #3) were evaluated by coercivity spectrum analysis using IRMUNMIX (Heslop et al., 2002) and IRM-CLG 1.0 (Kruiver et al., 2001) to determine the coercivities of the magnetic minerals (Heslop et al., 2004). The specimens were subjected to AF demagnetization up to 120 mT, and an IRM was imparted three orthogonal IRMs (120 mT, 500 mT, 2500 mT). Finally, the specimens were demagnetized thermally (Lowrie, 1980).

Strontium isotope analysis was utilized to determine if externally derived fluids altered the core. Analysis of 22 representative samples was performed at the University of Texas Thermal Ionization Mass Spectrometer laboratory. All samples were Osagean, with the exception of the deepest sample in the A. F. Severin core, a dark gray finely laminated argillaceous carbonate, which was assigned to the Kinderhookian. Samples were selected from all cores and included a mixture of matrix and fracture-filling calcite, dolomite in the matrix, calcite matrix, as well as some bulk samples. The samples were subjected to dissolution in acetic acid (4% dilution), and Sr was isolated by ion exchange chromatography. The mean value of NBS 987 for the experiment is 0.710275 ± .0000065 (2σ = 0.000013, n = 25). The 87Sr/86Sr values were normalized to NBS 987 = 0.71014 and compared to the coeval seawater value curve for the Mississippian (McArthur et al., 2001).

RESULTS AND INTERPRETATIONS

Paleomagnetic Analysis

Representative specimens from three cores were chosen to test if LTD treatment was necessary. Two liquid nitrogen treatments performed on 10 specimens from the A. F. Severin core increased intensity measurements by an average of 1.9% over NRM intensity. Treated and untreated specimens from the same core plug produced similar results, and the LTD treatment did not significantly improve the likelihood of identifying a ChRM. Samples from the Richard Chelf #1 and Tubbs #3 cores were cut into two specimens with one subjected to normal thermal demagnetization and the other specimen subjected to thermal demagnetization (10 steps from 40°C to 220°C) to isolate the VRM followed by two treatments of LTD. The LTD resulted in an average intensity drop of 2.6% for the Richard Chelf #1 core and 2.0% for the Tubbs #3 core. LTD treatment did not significantly improve decay or cause a significant change in the intensity of specimens. Based on these results, LTD was not performed for most specimens in this study.

Thermal demagnetization of many specimens removed a component at low temperatures (from the NRM −140°C to 100°C−280°C) that has steep positive inclinations (Figure 2A–F). The average inclination of this component (A. F. Severin = 64.6° [n = 50]; Oldenberg #1−16 = 61.2° [n = 27]; Sutton #1−16 = 67.4° [n = 12]; Richard Chelf #1 = 63.0° [n = 36]; Tubbs #3 = 58.9° [n = 24]) is close to the inclination of the modern field at the study location (64°), and this component is interpreted as a VRM.

Figure 2.

Representative Zijderveld diagrams. (A) A. F. Severin core (specimen AFS32c). Note VRM removed from the NRM −180°C and the ChRM between 280°C and 440°C. (B) A. F. Severin core (specimen AFS112b). Note removal of VRM from NRM −300°C and the ChRM from 300°C–400°C. (C) Oldenberg #1–16 specimen with the ChRM component removed between 200°C and 400°C. (D) Sutton #1–16 ChRM component removed between 200°C and 480°C. The first two steps before the NRM are LTD steps. (E) Richard Chelf #1 specimen with the ChRM component removed between 300°C and 480°C. The VRM component is from 100°C to 280°C (NRM removed). (F) Tubbs #3 ChRM specimen (NRM removed) with the ChRM removed between 260°C and 420°C. Squares: horizontal component; circles: vertical component

Figure 2.

Representative Zijderveld diagrams. (A) A. F. Severin core (specimen AFS32c). Note VRM removed from the NRM −180°C and the ChRM between 280°C and 440°C. (B) A. F. Severin core (specimen AFS112b). Note removal of VRM from NRM −300°C and the ChRM from 300°C–400°C. (C) Oldenberg #1–16 specimen with the ChRM component removed between 200°C and 400°C. (D) Sutton #1–16 ChRM component removed between 200°C and 480°C. The first two steps before the NRM are LTD steps. (E) Richard Chelf #1 specimen with the ChRM component removed between 300°C and 480°C. The VRM component is from 100°C to 280°C (NRM removed). (F) Tubbs #3 ChRM specimen (NRM removed) with the ChRM removed between 260°C and 420°C. Squares: horizontal component; circles: vertical component

At higher demagnetization temperatures, up to 440°C–500°C, a ChRM is removed in many specimens (Figure 2A–F). For those specimens treated with LTD that contained the ChRM, the maximum unblocking temperature was 500°C. Some specimens contained the VRM only, some contained the VRM and the ChRM, and some only the ChRM. A total of 620 specimens were processed from the five cores. Two hundred thirty-nine specimens contained the ChRM with a MAD of less than 15°. The specimen directions for the ChRM from the different cores have streaked declinations and shallow inclinations (Figure 3A, C, E). Most of these specimens without the ChRM came from the A. F. Severin core, which was the thickest core with 390 specimens processed and 106 with the ChRM (Figure 3A). Of the other cores, 52/89 of specimens from the Richard Chelf #1 (Figure 3C), 32/45 of specimens from the Tubbs #3, 37/47 of specimens from the Oldenberg #1–16 (Figure 3E), and 12/49 of specimens from the Sutton #1–16 had the ChRM. The specimens without the ChRM do not display stable linear decay or have MADs of greater than 15°. Specimens lacking the VRMs or ChRMs occur throughout every facies, and from specimens collected from the same sample plug as a specimen with the ChRM. Specimens were collected around some mineralized fractures. The ChRM is found in limestone with fractures as well as limestone without them. Because the fractures are not common or very thick, this contact test is considered inconclusive.

Figure 3.

Equal area projections of noncorrected and corrected (VRM [gray] method with VRMs rotated to 3.7° and the ChRMs rotated by the same amount) ChRM directions from three cores. (A) Uncorrected ChRM directions from the A. F. Severin core. (B) Declination corrected ChRMs from the A. F. Severin core. (C) Uncorrected ChRM directions from the Richard Chelf core. (D) Declination corrected ChRMs from the Richard Chelf core. (E) Uncorrected ChRM directions from the Oldenburg core. (F) Declination corrected ChRMs from the Oldenburg core. Note that the uncorrected directions for all cores have shallow inclinations and the corrected directions for all cores show a streak of declinations.

Figure 3.

Equal area projections of noncorrected and corrected (VRM [gray] method with VRMs rotated to 3.7° and the ChRMs rotated by the same amount) ChRM directions from three cores. (A) Uncorrected ChRM directions from the A. F. Severin core. (B) Declination corrected ChRMs from the A. F. Severin core. (C) Uncorrected ChRM directions from the Richard Chelf core. (D) Declination corrected ChRMs from the Richard Chelf core. (E) Uncorrected ChRM directions from the Oldenburg core. (F) Declination corrected ChRMs from the Oldenburg core. Note that the uncorrected directions for all cores have shallow inclinations and the corrected directions for all cores show a streak of declinations.

The VRM is assumed to be in the same direction as the modern field for the study location so it can be used to orient the specimens (e.g., Van der Voo and Watts, 1978; Suk et al., 1993) The VRMs were rotated to the modern direction (~3.7°) for the study location for each core, and the corresponding ChRMs were rotated by the same amount (Figure 3B, D, F). As can be seen from the results from the A. F. Severin core, the correction for the VRM produces a streak of declinations from 10° to 300° with shallow inclinations (Figure 3B). Richard Chelf #1 and Shell Oldenberg #1–16 (Figure 3D, F) as well as the Tubbs #3 (streak between 120.9° and 235.8° declination with shallow inclinations) produced similar results. There are only four specimens that contained the VRM and ChRM in the Sutton #1–16. These results indicate that the VRM method was not successful in orienting the cores. To gain a better understanding of the streaked data, the specimen directions were compared to depths, lithologies, and magnetic intensities to determine if any correlations existed but no correlations were found.

Although the average inclinations for the VRM in the cores (58.9°–67.4°) are close to the inclination of the modern field at the study location (64°), the possibility of a DIRM was tested by comparing possible differences between interior and exterior specimens with respect to the core. Drilling-induced magnetizations reside primarily in multidomain magnetite (Jackson and Van der Voo, 1985) and the magnetic intensities are commonly higher in the outer parts of the core (Wall and Worm, 2001) and have anomalously steep inclinations of 70°–90° (Burmester, 1977). Comparing VRM directions from inner and outer specimens shows no pattern in variation; the NRM intensities were not higher in the outer specimens. This, along with inclinations below 70°, suggests that a DIRM is not present.

Because cores are unoriented, a pole position could not be calculated and compared to the APWP to determine the time of remanence acquisition. The cores are all vertical, and the inclinations can be compared with expected inclinations to determine the magnetization age (Figure 4). The mean inclinations (inclination only; McFadden and Reid, 1982) are −1.6° for A. F. Severin, −2.5° for Oldenberg #1–16, −10.0° for Sutton #1–16, −5.3° for Richard Chelf #1 core, and 1.3° for Tubbs #3 core (Table 1). The mean for the ChRM for all specimens is −2.6° and the mean for the five cores is −3.6° (Table 1). The mean inclinations indicate that a remagnetization event occurred during the Permian (~310–290 Ma; Figure 4).

Figure 4.

Plot of expected inclination (in degrees) versus age (Ma) in the study area. The dashed line is the age of the Mississippian rocks. The thick filled line represents the average inclination (−2.6°) combined from all cores with the width of the line representing the α95 of 1.5°. Inclinations (thin black line) calculated from Torsvik et al. (2012).

Figure 4.

Plot of expected inclination (in degrees) versus age (Ma) in the study area. The dashed line is the age of the Mississippian rocks. The thick filled line represents the average inclination (−2.6°) combined from all cores with the width of the line representing the α95 of 1.5°. Inclinations (thin black line) calculated from Torsvik et al. (2012).

Table 1.

Characteristic remanent magnetization (ChRM) data—inclination only*.

CORENINCLINATIONα95k
A. F. Severin106−1.6°1.7°44.1
Oldeneberg #1–1637−2.5°3.3°35.0
Sutton #1–1612−10.0°−7.9°21.9
Richard Chelf #152−5.3°−2.4°44.5
Tubbs #3321.3°5.6°14.1
All specimen ChRMs239−2.6°1.5°30.7
Mean of cores5−3.6°5.6°180.8
CORENINCLINATIONα95k
A. F. Severin106−1.6°1.7°44.1
Oldeneberg #1–1637−2.5°3.3°35.0
Sutton #1–1612−10.0°−7.9°21.9
Richard Chelf #152−5.3°−2.4°44.5
Tubbs #3321.3°5.6°14.1
All specimen ChRMs239−2.6°1.5°30.7
Mean of cores5−3.6°5.6°180.8
*

N is the number of specimens with the ChRM; inclination is the “inclination only” calculated following McFadden and Reid (1982); α95 is the 95% cone of confidence; k is Fisher’s (1953) best estimate of the precision parameter.

Rock Magnetic Results

The magnetization in most specimens is interpreted to reside in magnetite because the ChRM in the majority of specimens decays above 440°C but below 580°C. Thirteen of the specimens with shallow inclinations have unblocking temperatures of 320°C –340°C, suggesting pyrrhotite may carry a magnetization in some specimens.

Representative IRM acquisition curves (Figure 5A–C) show that a low coercivity phases, such as magnetite and pyrrhotite, are the dominant magnetic minerals present in the cores because of saturation by 300–500 mT. The IRM acquisition data were analyzed using coercivity spectrum analysis (Kruiver et al., 2001) to determine the magnetic mineralogy in the A. F. Severin, Richard Chelf #1, and Tubbs #3 cores (Table 2). The primary component based on the IRM analysis is interpreted to be magnetite (Table 2) with the next most abundant component being pyrrhotite (Table 2). For example, in the AF Severin core, the primary component is interpreted to be magnetite (42.7 – 57.41 mT; Peters and Dekkers, 2003) and it contributes 48%–87% of the total IRM. Additional contributions to the total IRM are made by pyrrhotite; hematite and goethite are present but they are not considered to carry stable remanent magnetizations.

Table 2.

Coercivity spectrum analysis with interpreted magnetic mineral contributions.

SAMPLE (Depth ft)COMP.CONTRI. %SIRM A/mB1/2 mTDP mTMAGNETIC Mineralogy
RC18b164.43.63E-0439.80.30Magnetite
5397.8228.91.63E-0481.30.34Pyrrhotite
 36.73.67E-05229.10.42Hematite
RC27b164.33.45E-0440.70.35Magnetite
5414.85235.01.88E-0485.10.27Pyrrhotite
 30.60.44E-051349.00.13Goethite
RC30b153.26.61E-0435.50.33Magnetite
5419.35222.42.78E-0463.10.14Pyrrhotite
 324.53.04E-04128.80.41Hematite
RC54b110.03.29E-0510.20.11Magnetite
5465.0278.82.59E-0450.10.27Pyrrhotite
 311.13.66E-05151.40.42Hematite
RC68b159.82.96E-0435.50.3Magnetite
5516.9236.51.81E-0470.80.28Pyrrhotite
 33.72.83E-05354.80.48Hematite
RC82b122.02.68E-0455.00.16Magnetite
5550.95258.07.07E-0445.70.40Pyrrhotite
 320.02.43E-042570.40.26Goethite
T23c12.77.00E-0612.32.66Magnetite
5072.4279.32.06E-0446.80.30Pyrrhotite
 318.04.67E-05125.90.40Hematite
T39c118.63.56E-0458.90.13Magnetite
5119.8264.11.23-0353.70.38Pyrrhotite
 317.43.33E-042570.40.30Goethite
 115.14.39E-0514.10.2Magnetite
T42b271.02.07-0450.10.23Pyrrhotite
5126.6313.94.05E-05144.50.40Hematite
AFS17d147.84.40E-0548.190.29Magnetite
5949.8230.42.80E-0590.650.28Pyrrhotite
 321.92.01E-05544.500.26Hematite
AFS37a15.121.21E-0514.660.09Titanomagnetite
6054263.61.49E-0447.320.23Magnetite
 325.25.89E-05107.150.22Pyrrhotite
 46.11.42E-05249.460.26Hematite
AFS52b156.84.10E-0442.660.29Magnetite
6141.8235.42.56E-0488.810.25Pyrrhotite
 37.85.65E-05275.420.50Hematite
AFS64d159.66.00E-0545.710.28Magnetite
6191.1239.13.94E-05101.620.32Pyrrhotite
 31.31.31E-061570.360.12Goethite
AFS86d155.45.12E-0451.640.44Magnetite
6274.25237.33.45E-0463.100.23Pyrrhotite
 37.36.78E-05130.920.45Hematite
AFS94c12.61.65E-0512.300.03Titanomagnetite
6316.75287.35.48E-0457.410.28Magnetite
 310.06.30E-05263.030.45Hematite
AFS96a153.73.40E-0444.670.46Magnetite
6322.9240.02.53E-0460.260.22Pyrrhotite
 36.34.02E-05144.540.21Hematite
AFS100a158.11.12E-0446.770.34Magnetite
6342.4233.96.54E-0575.860.34Pyrrhotite
 38.01.55E-055495.410.58Goethite
AFS117b154.83.92E-0452.480.46Magnetite
6422.5236.42.60E-0464.570.23Pyrrhotite
 38.86.29E-05134.900.45Hematite
AFS121b181.52.64E-0463.100.32Pyrrhotite
6438.7213.94.50E-05549.540.28Hematite
 34.61.49E-051778.280.09Goethite
SAMPLE (Depth ft)COMP.CONTRI. %SIRM A/mB1/2 mTDP mTMAGNETIC Mineralogy
RC18b164.43.63E-0439.80.30Magnetite
5397.8228.91.63E-0481.30.34Pyrrhotite
 36.73.67E-05229.10.42Hematite
RC27b164.33.45E-0440.70.35Magnetite
5414.85235.01.88E-0485.10.27Pyrrhotite
 30.60.44E-051349.00.13Goethite
RC30b153.26.61E-0435.50.33Magnetite
5419.35222.42.78E-0463.10.14Pyrrhotite
 324.53.04E-04128.80.41Hematite
RC54b110.03.29E-0510.20.11Magnetite
5465.0278.82.59E-0450.10.27Pyrrhotite
 311.13.66E-05151.40.42Hematite
RC68b159.82.96E-0435.50.3Magnetite
5516.9236.51.81E-0470.80.28Pyrrhotite
 33.72.83E-05354.80.48Hematite
RC82b122.02.68E-0455.00.16Magnetite
5550.95258.07.07E-0445.70.40Pyrrhotite
 320.02.43E-042570.40.26Goethite
T23c12.77.00E-0612.32.66Magnetite
5072.4279.32.06E-0446.80.30Pyrrhotite
 318.04.67E-05125.90.40Hematite
T39c118.63.56E-0458.90.13Magnetite
5119.8264.11.23-0353.70.38Pyrrhotite
 317.43.33E-042570.40.30Goethite
 115.14.39E-0514.10.2Magnetite
T42b271.02.07-0450.10.23Pyrrhotite
5126.6313.94.05E-05144.50.40Hematite
AFS17d147.84.40E-0548.190.29Magnetite
5949.8230.42.80E-0590.650.28Pyrrhotite
 321.92.01E-05544.500.26Hematite
AFS37a15.121.21E-0514.660.09Titanomagnetite
6054263.61.49E-0447.320.23Magnetite
 325.25.89E-05107.150.22Pyrrhotite
 46.11.42E-05249.460.26Hematite
AFS52b156.84.10E-0442.660.29Magnetite
6141.8235.42.56E-0488.810.25Pyrrhotite
 37.85.65E-05275.420.50Hematite
AFS64d159.66.00E-0545.710.28Magnetite
6191.1239.13.94E-05101.620.32Pyrrhotite
 31.31.31E-061570.360.12Goethite
AFS86d155.45.12E-0451.640.44Magnetite
6274.25237.33.45E-0463.100.23Pyrrhotite
 37.36.78E-05130.920.45Hematite
AFS94c12.61.65E-0512.300.03Titanomagnetite
6316.75287.35.48E-0457.410.28Magnetite
 310.06.30E-05263.030.45Hematite
AFS96a153.73.40E-0444.670.46Magnetite
6322.9240.02.53E-0460.260.22Pyrrhotite
 36.34.02E-05144.540.21Hematite
AFS100a158.11.12E-0446.770.34Magnetite
6342.4233.96.54E-0575.860.34Pyrrhotite
 38.01.55E-055495.410.58Goethite
AFS117b154.83.92E-0452.480.46Magnetite
6422.5236.42.60E-0464.570.23Pyrrhotite
 38.86.29E-05134.900.45Hematite
AFS121b181.52.64E-0463.100.32Pyrrhotite
6438.7213.94.50E-05549.540.28Hematite
 34.61.49E-051778.280.09Goethite

Comp. is the order the components; Contri. is the relative contribution of each component; SIRM is the saturation isothermal remanent magnetization; B(1/2) is the field where half of the SIRM is reached; DP, dispersion parameter, is one standard deviation of the logarithmic distribution.

Figure 5.

Representative IRM acquisition curves for the (A) A. F. Severin, (B) Richard Chelf, and (C) Sutton cores. The curves show saturation by 300–500 mT, suggesting the presence of a low coercivity mineral or minerals.

Figure 5.

Representative IRM acquisition curves for the (A) A. F. Severin, (B) Richard Chelf, and (C) Sutton cores. The curves show saturation by 300–500 mT, suggesting the presence of a low coercivity mineral or minerals.

Thermal decay of representative specimens with a triaxial IRM shows that the low coercivity component (120 mT) has the highest intensities (Figure 6A–C). Most specimens (Figure 6A, C) are characterized by gradual decay of the low coercivity component up to 420°C–500°C, suggesting the presence of magnetite. Some specimens are characterized by steep decay around 300°C –320°C (Figure 6B) and then more gradual decay to 420°C–500°C. The change in slope near 320°C is consistent with unblocking temperatures for pyrrhotite (325°C), whereas decay to higher temperatures indicates magnetite.

Figure 6.

Representative thermal decay curves of a triaxial IRM for (A) A. F. Severin specimen, (B) Richard Chelf #1 specimen, and (C) Sutton specimen. The change in slope near 320°C in some specimens is consistent with unblocking temperatures for pyrrhotite (325°C) whereas decay to higher temperatures indicates magnetite. Note that the 2,500 mT curve in (A) has a higher intensity than the 500 mT curve compared to the other specimens. This suggests that the sample may contain more hematite but it is not interpreted to carry a stable magnetization.

Figure 6.

Representative thermal decay curves of a triaxial IRM for (A) A. F. Severin specimen, (B) Richard Chelf #1 specimen, and (C) Sutton specimen. The change in slope near 320°C in some specimens is consistent with unblocking temperatures for pyrrhotite (325°C) whereas decay to higher temperatures indicates magnetite. Note that the 2,500 mT curve in (A) has a higher intensity than the 500 mT curve compared to the other specimens. This suggests that the sample may contain more hematite but it is not interpreted to carry a stable magnetization.

Facies and Paragenesis

Five facies were identified in the cores: chert breccia (Figure 7A), shaly mudstone, cherty wackestone or packstone, peloidal packstone, and glauconitic sandstone. These facies are similar to what other workers have identified (e.g., Mazzullo et al., 2009, 2011; Young, 2010) and will not be described in this paper. Most facies are present in all cores with the exception of chert breccia, which is not present in the Shell Sutton #1–16 and Shell Oldenberg #1–16 cores, and the glauconitic sandstone facies is only found in the A. F. Severin core. In the middle of the Richard Chelf #1 core, there is a 0.6 m (2 ft) chert breccia (5453 ft [1162 m] to 5455 ft [1163 m]) that could be a fracture-fill (Figure 7B). This breccia may have been caused by subaerial karsting and collapse, hydrothermal alteration and dissolution, and/or fault tectonism. It could have been a conduit for fluid flow (e.g., Elebiju et al., 2011).

Figure 7.

Photomicrographs and pictures of diagenetic features in the cores. (A) Photomicrograph of a brecciated interval in the Richard Chelf core. (B) Core slab photograph of the brecciated zone in the Richard Chelf core that is hypothesized to be a filled fracture. (C) Multiple generations of chalcedony followed by megaquartz in a fracture with a chert matrix in the Sutton core (x-polarized light). (D) Baroque dolomite filling a vug and fracture from the A. F. Severin core (x-polarized light). (E) Scanning electron photomicrograph (backscatter) of sphalerite (sl) from the Severin core (cal = calcite; py = pyrite).

Figure 7.

Photomicrographs and pictures of diagenetic features in the cores. (A) Photomicrograph of a brecciated interval in the Richard Chelf core. (B) Core slab photograph of the brecciated zone in the Richard Chelf core that is hypothesized to be a filled fracture. (C) Multiple generations of chalcedony followed by megaquartz in a fracture with a chert matrix in the Sutton core (x-polarized light). (D) Baroque dolomite filling a vug and fracture from the A. F. Severin core (x-polarized light). (E) Scanning electron photomicrograph (backscatter) of sphalerite (sl) from the Severin core (cal = calcite; py = pyrite).

The diagenetic events can be subdivided into several stages. The early period of diagenesis includes micritization, initial calcite cementation, and early dissolution of some carbonate material, a first stage of silica (chert and chalcedony) and dolomite cementation. Chert and chalcedony are both cut by early blocky (coarsely crystalline) calcite-filled fractures. Middle diagenesis is characterized by brecciation (Figure 7A); deformed fractures filled with blocky calcite, silica, and carbonate dissolution; a second generation of chalcedony and chert (Figure 7C); and replacement by dolomite. Silica dissolution, as well as brecciation, has been interpreted as occurring during subaerial exposure in the Mississippian (e.g., Montgomery et al., 1998; Rogers, 2001; Franseen, 2006; Mazzullo et al., 2009, 2011). Late-stage diagenesis includes angular fractures and vugs filled with baroque dolomite in the A. F. Severin core (Figure 7D). Sphalerite is found in a few thin sections, generally associated with pyrite (Figure 7E). It could be hydrothermal in origin or it may have precipitated from internal fluids. Hydrocarbons have also been found in all the cores with the exception of the A. F. Severin core. There is some variability in the paragenesis among the cores, but there are more similarities than differences.

Geochemistry

The 87Sr/86Sr data from the cores were compared with the seawater values for the Osagean and Kinderhookian series (McArthur et al., 2001). Most samples plot from slightly above coeval to slightly depleted when compared to the Osagean portion of the Mississippian strontium seawater curve (Figure 8). One sample of chert with dolomite from the Richard Chelf core has a slightly elevated 87Sr/86Sr value. The value of the clay-rich Kinderhookian sample is elevated compared to the Sr seawater curve (Figure 8). There was no discernable pattern between the values and different types of samples analyzed in the other samples.

Figure 8.

The 87Sr/86Sr results plotted relative to the Sr values for Mississippian seawater. The middle curve is the mean value with the errors shown by the upper and lower curves (McArthur et al., 2001). Severin: ellipses; Richard Chelf #1: stars; Tubbs #3: circles; Oldenberg #1–16: polygons; Sutton #1–6: squares (modified from McArthur et al., 2001) (C = Chesterian, K = Kinderhookian, M = Meramecian, and O = Osagean).

Figure 8.

The 87Sr/86Sr results plotted relative to the Sr values for Mississippian seawater. The middle curve is the mean value with the errors shown by the upper and lower curves (McArthur et al., 2001). Severin: ellipses; Richard Chelf #1: stars; Tubbs #3: circles; Oldenberg #1–16: polygons; Sutton #1–6: squares (modified from McArthur et al., 2001) (C = Chesterian, K = Kinderhookian, M = Meramecian, and O = Osagean).

DISCUSSION

The majority of specimens have maximum unblocking temperatures above 360°C and below 580°C, the Curie temperature for magnetite. This suggests that the ChRM is carried in magnetite and this interpretation is consistent with the rock magnetic results. Thirteen specimens have maximum unblocking temperatures near 325°C, the Curie temperature for pyrrhotite; this combined with IRM results indicate the ChRM in some specimens may reside in pyrrhotite. Hematite and goethite are also present based on the rock magnetic results but they are not interpreted to carry a stable magnetization.

The VRM orienting method produced a set of streaked ChRM directions in all five cores and was not successful in orienting the cores. One possible explanation is that VRM may be contaminated by an unresolved DIRM but the VRM inclinations in the cores are all below 70° and there is no difference in the magnetization between outer and inner cores. This suggests that a DIRM is not an explanation for failure of the VRM orienting method. Another possibility is that the VRM is not stable or well defined enough to orient in these cores. Determining the reason the VRM was not able to be used to orient the cores is an unresolved issue that is beyond the scope of this study, but should be addressed in future research.

The Permian age for the ChRM indicates it is a secondary magnetization, either a thermoviscous remanent magnetization (TVRM) or a CRM. Vitrinite reflectance values for the Woodford Shale in the study are less than 1.0 (Cardott and Lambert, 1995), which suggests burial temperatures of approximately 100°C –150°C (Barker and Pawlewicz, 1994). Other paleotemperature estimates are 90°C for the Mississippian in southeast Kansas (Förster et al., 1998) and 93°C –144°C from fluid inclusions in Kansas (Goldstein and King, 2014). Assuming a maximum burial temperature near the upper range of these values (140°C) and using the temperature relaxation time curves (Dunlop et al., 2000), the maximum unblocking temperatures near 500°C are too high for the magnetite ChRM to be a TVRM. The relaxation time–temperature curves for single domain magnetite (Dunlop et al., 2000) can be used because specimens were treated with LTD. Similarly, a TVRM origin for a pyrrhotite ChRM can be ruled out, based on maximum unblocking temperatures of 320°C and the relaxation time–temperature curves for pyrrhotite (Dunlop et al., 2000). Therefore, the ChRMs residing in magnetite, and perhaps pyrrhotite, are interpreted as CRMs.

There are several possible chemical remagnetization mechanisms that could account for the CRM in magnetite (e.g., Elmore et al., 2012). These include fluid-related and burial processes such as smectite to illite conversion (Katz et al., 2000) or maturation of organic matter (Banerjee et al., 1997; Blumstein et al., 2004). The carbonates do not contain abundant clay, suggesting that smectite to illite conversion is not a likely mechanism for magnetite authigenesis. In addition, shale present in the shale and mudstone facies is the most clay-rich material, but no samples from the shale contained the CRM. The Mississippian Limestone contains some organic matter, and some intervals could be a source rock (Watney, 2014). Maturation of organic matter, therefore, is a possible chemical remagnetization mechanism.

Migration of hydrocarbons is another potential process that can cause magnetite authigenesis and CRM acquisition (e.g., Elmore et al., 1993). The Woodford Shale is the likely source rock for the Mississippian strata in northern Oklahoma and the unit reached peak burial during the Permian (Förster et al., 1998), although it generated hydrocarbons as early as 335 Ma (Gaswirth and Higley, 2013). A Permian timing of hydrocarbon migration is consistent with the age of the CRM but hydrocarbons are not abundant in the cores.

There is strong evidence that dissolution by meteoric fluids at unconformities resulted in the creation of porosity in the Mississippi limestone (e.g., Rogers et al., 1995; Montgomery et al., 1998; Watney et al., 2001; Franseen, 2006; Mazzullo et al., 2009). Recent studies (e.g., Goldstein and King, 2014; Ramaker et al., 2014) suggest that hydrothermal fluid flow may have enhanced the porosity in the reservoirs. Goldstein and King (2014) propose a three-stage hydrothermal system that affected porosity development. The first stage had relatively low salinities and was interpreted to be a Late Pennsylvanian or early Permian event. The second stage was interpreted to be Permian in age and was characterized by higher salinities and precipitation of baroque dolomite. This event was interpreted to have been affected by Ouachita tectonism. The final stage was interpreted to be a result of more localized fluid flow likely after Ouachita deformation and possibly related to Laramide reactivation.

A Permian age for CRM acquisition is consistent with the second stage of the hydrothermal system proposed by Goldstein and King (2014). The presence of baroque dolomite and sphalerite is also consistent with this interpretation. In addition, MVT deposits in the Tri-State Mineral District formed from hydrothermal fluids (Coveney, 1992) and the CRM timing in the Mississippian limestone is coincident with paleomagnetic dates (Pan et al., 1990; Symons et al., 2005) as well as radiometric dates (Brannon et al., 1996) for the MVT mineralization. The 87Sr/86Sr data from this study show that most samples are close to the coeval seawater values (McArthur et al., 2001) although the results from a few samples have higher values than the coeval Sr seawater curve (Figure 8), which may suggest alteration by evolved fluids.

Based on the Permian timing and the presence of minerals of possible hydrothermal origin, the CRM is interpreted to have been caused by hydrothermal fluids. Other mechanisms such as alteration by hydrocarbons or maturation of organic matter, however, cannot be ruled out as a remagnetization mechanism. We note that oil migration may have been associated the regional hydrothermal systems (e.g., stage 2, Goldstein and King, 2014).

CONCLUSIONS

The paleomagnetic data from five cores show that the Mississippian limestone in northern Oklahoma contains a Permian CRM that resides in magnetite. Rock magnetic results are consistent with the magnetite interpretation. The cores were unoriented, and an attempt was made to orient them using the VRM method. This approach resulted in a streaked distribution of declinations and shallow inclinations and was not successful. As a result, the inclinations of the CRM in the five cores were compared to the expected inclinations for the study area, which indicates remanence acquisition in the Permian (~310–290 Ma). Petrographic analysis indicates a complex paragenetic sequence, which includes dolomitization, brecciation, and precipitation of quartz, calcite, sphalerite, pyrite, and baroque dolomite.

The age of the CRM is consistent with dates for mineralization in the nearby Tri-State MVT deposit and for a hypothesized regional Permian hydrothermal event. The age of the CRM and the presence of sphalerite and baroque dolomite suggest that the CRM was acquired via external hydrothermal fluids in the Permian. The results are consistent with the hypothesis that a hydrothermal system affected the rocks and probably enhanced the reservoir quality.

ACKNOWLEDGMENTS

The authors thank Devon Energy for providing support, Melissa Bennet for help processing the specimens, and Alyssa Wickard for help drafting some of the figures.

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Figures & Tables

Figure 1.

Location map showing the cores sampled in this study (modified from Oklahoma Geological Survey map).

Figure 1.

Location map showing the cores sampled in this study (modified from Oklahoma Geological Survey map).

Figure 2.

Representative Zijderveld diagrams. (A) A. F. Severin core (specimen AFS32c). Note VRM removed from the NRM −180°C and the ChRM between 280°C and 440°C. (B) A. F. Severin core (specimen AFS112b). Note removal of VRM from NRM −300°C and the ChRM from 300°C–400°C. (C) Oldenberg #1–16 specimen with the ChRM component removed between 200°C and 400°C. (D) Sutton #1–16 ChRM component removed between 200°C and 480°C. The first two steps before the NRM are LTD steps. (E) Richard Chelf #1 specimen with the ChRM component removed between 300°C and 480°C. The VRM component is from 100°C to 280°C (NRM removed). (F) Tubbs #3 ChRM specimen (NRM removed) with the ChRM removed between 260°C and 420°C. Squares: horizontal component; circles: vertical component

Figure 2.

Representative Zijderveld diagrams. (A) A. F. Severin core (specimen AFS32c). Note VRM removed from the NRM −180°C and the ChRM between 280°C and 440°C. (B) A. F. Severin core (specimen AFS112b). Note removal of VRM from NRM −300°C and the ChRM from 300°C–400°C. (C) Oldenberg #1–16 specimen with the ChRM component removed between 200°C and 400°C. (D) Sutton #1–16 ChRM component removed between 200°C and 480°C. The first two steps before the NRM are LTD steps. (E) Richard Chelf #1 specimen with the ChRM component removed between 300°C and 480°C. The VRM component is from 100°C to 280°C (NRM removed). (F) Tubbs #3 ChRM specimen (NRM removed) with the ChRM removed between 260°C and 420°C. Squares: horizontal component; circles: vertical component

Figure 3.

Equal area projections of noncorrected and corrected (VRM [gray] method with VRMs rotated to 3.7° and the ChRMs rotated by the same amount) ChRM directions from three cores. (A) Uncorrected ChRM directions from the A. F. Severin core. (B) Declination corrected ChRMs from the A. F. Severin core. (C) Uncorrected ChRM directions from the Richard Chelf core. (D) Declination corrected ChRMs from the Richard Chelf core. (E) Uncorrected ChRM directions from the Oldenburg core. (F) Declination corrected ChRMs from the Oldenburg core. Note that the uncorrected directions for all cores have shallow inclinations and the corrected directions for all cores show a streak of declinations.

Figure 3.

Equal area projections of noncorrected and corrected (VRM [gray] method with VRMs rotated to 3.7° and the ChRMs rotated by the same amount) ChRM directions from three cores. (A) Uncorrected ChRM directions from the A. F. Severin core. (B) Declination corrected ChRMs from the A. F. Severin core. (C) Uncorrected ChRM directions from the Richard Chelf core. (D) Declination corrected ChRMs from the Richard Chelf core. (E) Uncorrected ChRM directions from the Oldenburg core. (F) Declination corrected ChRMs from the Oldenburg core. Note that the uncorrected directions for all cores have shallow inclinations and the corrected directions for all cores show a streak of declinations.

Figure 4.

Plot of expected inclination (in degrees) versus age (Ma) in the study area. The dashed line is the age of the Mississippian rocks. The thick filled line represents the average inclination (−2.6°) combined from all cores with the width of the line representing the α95 of 1.5°. Inclinations (thin black line) calculated from Torsvik et al. (2012).

Figure 4.

Plot of expected inclination (in degrees) versus age (Ma) in the study area. The dashed line is the age of the Mississippian rocks. The thick filled line represents the average inclination (−2.6°) combined from all cores with the width of the line representing the α95 of 1.5°. Inclinations (thin black line) calculated from Torsvik et al. (2012).

Figure 5.

Representative IRM acquisition curves for the (A) A. F. Severin, (B) Richard Chelf, and (C) Sutton cores. The curves show saturation by 300–500 mT, suggesting the presence of a low coercivity mineral or minerals.

Figure 5.

Representative IRM acquisition curves for the (A) A. F. Severin, (B) Richard Chelf, and (C) Sutton cores. The curves show saturation by 300–500 mT, suggesting the presence of a low coercivity mineral or minerals.

Figure 6.

Representative thermal decay curves of a triaxial IRM for (A) A. F. Severin specimen, (B) Richard Chelf #1 specimen, and (C) Sutton specimen. The change in slope near 320°C in some specimens is consistent with unblocking temperatures for pyrrhotite (325°C) whereas decay to higher temperatures indicates magnetite. Note that the 2,500 mT curve in (A) has a higher intensity than the 500 mT curve compared to the other specimens. This suggests that the sample may contain more hematite but it is not interpreted to carry a stable magnetization.

Figure 6.

Representative thermal decay curves of a triaxial IRM for (A) A. F. Severin specimen, (B) Richard Chelf #1 specimen, and (C) Sutton specimen. The change in slope near 320°C in some specimens is consistent with unblocking temperatures for pyrrhotite (325°C) whereas decay to higher temperatures indicates magnetite. Note that the 2,500 mT curve in (A) has a higher intensity than the 500 mT curve compared to the other specimens. This suggests that the sample may contain more hematite but it is not interpreted to carry a stable magnetization.

Figure 7.

Photomicrographs and pictures of diagenetic features in the cores. (A) Photomicrograph of a brecciated interval in the Richard Chelf core. (B) Core slab photograph of the brecciated zone in the Richard Chelf core that is hypothesized to be a filled fracture. (C) Multiple generations of chalcedony followed by megaquartz in a fracture with a chert matrix in the Sutton core (x-polarized light). (D) Baroque dolomite filling a vug and fracture from the A. F. Severin core (x-polarized light). (E) Scanning electron photomicrograph (backscatter) of sphalerite (sl) from the Severin core (cal = calcite; py = pyrite).

Figure 7.

Photomicrographs and pictures of diagenetic features in the cores. (A) Photomicrograph of a brecciated interval in the Richard Chelf core. (B) Core slab photograph of the brecciated zone in the Richard Chelf core that is hypothesized to be a filled fracture. (C) Multiple generations of chalcedony followed by megaquartz in a fracture with a chert matrix in the Sutton core (x-polarized light). (D) Baroque dolomite filling a vug and fracture from the A. F. Severin core (x-polarized light). (E) Scanning electron photomicrograph (backscatter) of sphalerite (sl) from the Severin core (cal = calcite; py = pyrite).

Figure 8.

The 87Sr/86Sr results plotted relative to the Sr values for Mississippian seawater. The middle curve is the mean value with the errors shown by the upper and lower curves (McArthur et al., 2001). Severin: ellipses; Richard Chelf #1: stars; Tubbs #3: circles; Oldenberg #1–16: polygons; Sutton #1–6: squares (modified from McArthur et al., 2001) (C = Chesterian, K = Kinderhookian, M = Meramecian, and O = Osagean).

Figure 8.

The 87Sr/86Sr results plotted relative to the Sr values for Mississippian seawater. The middle curve is the mean value with the errors shown by the upper and lower curves (McArthur et al., 2001). Severin: ellipses; Richard Chelf #1: stars; Tubbs #3: circles; Oldenberg #1–16: polygons; Sutton #1–6: squares (modified from McArthur et al., 2001) (C = Chesterian, K = Kinderhookian, M = Meramecian, and O = Osagean).

Table 1.

Characteristic remanent magnetization (ChRM) data—inclination only*.

CORENINCLINATIONα95k
A. F. Severin106−1.6°1.7°44.1
Oldeneberg #1–1637−2.5°3.3°35.0
Sutton #1–1612−10.0°−7.9°21.9
Richard Chelf #152−5.3°−2.4°44.5
Tubbs #3321.3°5.6°14.1
All specimen ChRMs239−2.6°1.5°30.7
Mean of cores5−3.6°5.6°180.8
CORENINCLINATIONα95k
A. F. Severin106−1.6°1.7°44.1
Oldeneberg #1–1637−2.5°3.3°35.0
Sutton #1–1612−10.0°−7.9°21.9
Richard Chelf #152−5.3°−2.4°44.5
Tubbs #3321.3°5.6°14.1
All specimen ChRMs239−2.6°1.5°30.7
Mean of cores5−3.6°5.6°180.8
*

N is the number of specimens with the ChRM; inclination is the “inclination only” calculated following McFadden and Reid (1982); α95 is the 95% cone of confidence; k is Fisher’s (1953) best estimate of the precision parameter.

Table 2.

Coercivity spectrum analysis with interpreted magnetic mineral contributions.

SAMPLE (Depth ft)COMP.CONTRI. %SIRM A/mB1/2 mTDP mTMAGNETIC Mineralogy
RC18b164.43.63E-0439.80.30Magnetite
5397.8228.91.63E-0481.30.34Pyrrhotite
 36.73.67E-05229.10.42Hematite
RC27b164.33.45E-0440.70.35Magnetite
5414.85235.01.88E-0485.10.27Pyrrhotite
 30.60.44E-051349.00.13Goethite
RC30b153.26.61E-0435.50.33Magnetite
5419.35222.42.78E-0463.10.14Pyrrhotite
 324.53.04E-04128.80.41Hematite
RC54b110.03.29E-0510.20.11Magnetite
5465.0278.82.59E-0450.10.27Pyrrhotite
 311.13.66E-05151.40.42Hematite
RC68b159.82.96E-0435.50.3Magnetite
5516.9236.51.81E-0470.80.28Pyrrhotite
 33.72.83E-05354.80.48Hematite
RC82b122.02.68E-0455.00.16Magnetite
5550.95258.07.07E-0445.70.40Pyrrhotite
 320.02.43E-042570.40.26Goethite
T23c12.77.00E-0612.32.66Magnetite
5072.4279.32.06E-0446.80.30Pyrrhotite
 318.04.67E-05125.90.40Hematite
T39c118.63.56E-0458.90.13Magnetite
5119.8264.11.23-0353.70.38Pyrrhotite
 317.43.33E-042570.40.30Goethite
 115.14.39E-0514.10.2Magnetite
T42b271.02.07-0450.10.23Pyrrhotite
5126.6313.94.05E-05144.50.40Hematite
AFS17d147.84.40E-0548.190.29Magnetite
5949.8230.42.80E-0590.650.28Pyrrhotite
 321.92.01E-05544.500.26Hematite
AFS37a15.121.21E-0514.660.09Titanomagnetite
6054263.61.49E-0447.320.23Magnetite
 325.25.89E-05107.150.22Pyrrhotite
 46.11.42E-05249.460.26Hematite
AFS52b156.84.10E-0442.660.29Magnetite
6141.8235.42.56E-0488.810.25Pyrrhotite
 37.85.65E-05275.420.50Hematite
AFS64d159.66.00E-0545.710.28Magnetite
6191.1239.13.94E-05101.620.32Pyrrhotite
 31.31.31E-061570.360.12Goethite
AFS86d155.45.12E-0451.640.44Magnetite
6274.25237.33.45E-0463.100.23Pyrrhotite
 37.36.78E-05130.920.45Hematite
AFS94c12.61.65E-0512.300.03Titanomagnetite
6316.75287.35.48E-0457.410.28Magnetite
 310.06.30E-05263.030.45Hematite
AFS96a153.73.40E-0444.670.46Magnetite
6322.9240.02.53E-0460.260.22Pyrrhotite
 36.34.02E-05144.540.21Hematite
AFS100a158.11.12E-0446.770.34Magnetite
6342.4233.96.54E-0575.860.34Pyrrhotite
 38.01.55E-055495.410.58Goethite
AFS117b154.83.92E-0452.480.46Magnetite
6422.5236.42.60E-0464.570.23Pyrrhotite
 38.86.29E-05134.900.45Hematite
AFS121b181.52.64E-0463.100.32Pyrrhotite
6438.7213.94.50E-05549.540.28Hematite
 34.61.49E-051778.280.09Goethite
SAMPLE (Depth ft)COMP.CONTRI. %SIRM A/mB1/2 mTDP mTMAGNETIC Mineralogy
RC18b164.43.63E-0439.80.30Magnetite
5397.8228.91.63E-0481.30.34Pyrrhotite
 36.73.67E-05229.10.42Hematite
RC27b164.33.45E-0440.70.35Magnetite
5414.85235.01.88E-0485.10.27Pyrrhotite
 30.60.44E-051349.00.13Goethite
RC30b153.26.61E-0435.50.33Magnetite
5419.35222.42.78E-0463.10.14Pyrrhotite
 324.53.04E-04128.80.41Hematite
RC54b110.03.29E-0510.20.11Magnetite
5465.0278.82.59E-0450.10.27Pyrrhotite
 311.13.66E-05151.40.42Hematite
RC68b159.82.96E-0435.50.3Magnetite
5516.9236.51.81E-0470.80.28Pyrrhotite
 33.72.83E-05354.80.48Hematite
RC82b122.02.68E-0455.00.16Magnetite
5550.95258.07.07E-0445.70.40Pyrrhotite
 320.02.43E-042570.40.26Goethite
T23c12.77.00E-0612.32.66Magnetite
5072.4279.32.06E-0446.80.30Pyrrhotite
 318.04.67E-05125.90.40Hematite
T39c118.63.56E-0458.90.13Magnetite
5119.8264.11.23-0353.70.38Pyrrhotite
 317.43.33E-042570.40.30Goethite
 115.14.39E-0514.10.2Magnetite
T42b271.02.07-0450.10.23Pyrrhotite
5126.6313.94.05E-05144.50.40Hematite
AFS17d147.84.40E-0548.190.29Magnetite
5949.8230.42.80E-0590.650.28Pyrrhotite
 321.92.01E-05544.500.26Hematite
AFS37a15.121.21E-0514.660.09Titanomagnetite
6054263.61.49E-0447.320.23Magnetite
 325.25.89E-05107.150.22Pyrrhotite
 46.11.42E-05249.460.26Hematite
AFS52b156.84.10E-0442.660.29Magnetite
6141.8235.42.56E-0488.810.25Pyrrhotite
 37.85.65E-05275.420.50Hematite
AFS64d159.66.00E-0545.710.28Magnetite
6191.1239.13.94E-05101.620.32Pyrrhotite
 31.31.31E-061570.360.12Goethite
AFS86d155.45.12E-0451.640.44Magnetite
6274.25237.33.45E-0463.100.23Pyrrhotite
 37.36.78E-05130.920.45Hematite
AFS94c12.61.65E-0512.300.03Titanomagnetite
6316.75287.35.48E-0457.410.28Magnetite
 310.06.30E-05263.030.45Hematite
AFS96a153.73.40E-0444.670.46Magnetite
6322.9240.02.53E-0460.260.22Pyrrhotite
 36.34.02E-05144.540.21Hematite
AFS100a158.11.12E-0446.770.34Magnetite
6342.4233.96.54E-0575.860.34Pyrrhotite
 38.01.55E-055495.410.58Goethite
AFS117b154.83.92E-0452.480.46Magnetite
6422.5236.42.60E-0464.570.23Pyrrhotite
 38.86.29E-05134.900.45Hematite
AFS121b181.52.64E-0463.100.32Pyrrhotite
6438.7213.94.50E-05549.540.28Hematite
 34.61.49E-051778.280.09Goethite

Comp. is the order the components; Contri. is the relative contribution of each component; SIRM is the saturation isothermal remanent magnetization; B(1/2) is the field where half of the SIRM is reached; DP, dispersion parameter, is one standard deviation of the logarithmic distribution.

Contents

GeoRef

References

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