- © 2012 UW Department of Geology and Geophysics
The Eureka Quartzite is a sheet-like quartzarenite up to 200-m thick that was deposited on the eastern shelf of the Cordilleran miogeocline from Canada to California. It is the only sandstone lithosome from the Middle Cambrian through Devonian succession in the Great Basin and is remarkable in its purity of detrital and authigenic quartz, scarcity of bedding, and heterogeneity of both grain packing and quartz cement abundance. Sand sources ranged from the Peace River Arch in Canada to the Transcontinental Arch in mid-continental North America. The Eureka represents a third-order regressive–transgressive stratigraphic sequence, although whether the regression formed in response to eustasy or epirogenic uplift of western North America is unresolved.
The near absence of detrital clay, body fossils, and subaerial features in addition to the presence of herringbone cross-beds indicate that the Eureka was deposited in intertidal and shallow subtidal environments except for minor eolian deposits. Bioturbation destroyed most primary stratification, although discrete burrows are rare. Textural features of quartz (superb roundness, bean grain shape, crescentic impact scars) indicate a prolonged episode of eolian abrasion prior to marine deposition.
Regionally the detrital composition is 99.5 percent detrital monocrystalline quartz, 0.5 percent K-feldspar and carbonate allochems, and a trace of heavy minerals. From 2 percent to 4 percent feldspar and carbonate allochems that were initially present have been replaced by quartz during burial. The chief authigenic phases are quartz overgrowths with minor calcite (now dolomite) and illite. Spheroidal and amoeboid calcite-cemented concretions up to 3 cm in diameter formed at shallow burial depths, but all carbonate in the concretions has been leached in outcrop. The heterogeneity of grain compaction and amount of quartz cement resulted in beds that range from semifriable to sedimentary quartzites in the same outcrop. Compaction by the combination of grain rearrangement, pressure dissolution, and grain fracturing generated anomalously low intergranular volumes that average 14 percent in Nevada and 21 percent in Utah.
The normal evolution of microquartz overgrowths (<10 μm) to meso and macroquartz overgrowths (>10 μm) during cementation was retarded; consequently, microquartz and mesoquartz cement (80 percent) dominate over macroquartz (20 percent), and they are the only cements in the least-cemented beds and laminations. Illite co-precipitated with microquartz and impeded quartz cementation by coating quartz crystal faces. Despite reaching temperatures >135° C for ∼100 million years, much of the Eureka, especially in Utah, remains incompletely cemented and retains porosity of ∼2 percent. The chief cause of cement heterogeneity appears to be authigenic illite abundance. Pressure dissolution of quartz at shale beds and clay drapes that formed stylolites was the most likely major source of silica for quartz cement.
Invasion by hydrocarbons and hydrogen sulfide (H2S) resulted in the reduction of iron in hematite grain coats on quartz grains, a remnant of their eolian dune formation, the bleaching of beds, and the generation of pyrite (now hematite). The abundance of iron that now resides in pyrite/hematite suggests that red beds were once widespread. Outcrop and near-outcrop processes generated Liesegang bands of iron oxide, desert varnish, “pockmarks” where carbonate cement in centimeter-scale concretions dissolved, hematite pseudomorphs of pyrite, and minor opal cement.
- quartz cement
- shallow marine
The Neoproterozoic breakup of the supercontinent Rodinia resulted in the development of a passive continental margin on the western side of the North American craton (Dickinson, 2004). Deposition on the eastern shelf of this Cordilleran miogeocline or foreland basin was nearly continuous from the Neoproterozoic through Late Devonian (Poole et al., 1992). By Late Devonian time, oceanic allochtons were thrust eastward across the miogeoclinal belt until Early Triassic time (Dickinson, 2004). In the area of the Great Basin in Nevada and Utah, deposition on the paleoshelf was dominated by carbonate and minor shale beds from Middle Cambrian through Devonian time, a span of more than 200 million years (Poole et al., 1992; Miall, 2010). During part of Middle and Late Ordovician time, however, well-rounded quartzarenite sand was deposited across a large part of the shelf from west-central Canada to California. The preserved area of this sand sheet in eastern California, Utah, Nevada, Idaho, British Columbia, and Alberta reaches 450,000 km2 (Ketner, 1968). Where exposed, the sandstone sheet varies regionally in thickness from 40 m to 200 m in the United States but reaches 520 m in the area of the Peace River Arch in Alberta, Canada (Ketner, 1968). The major part of this sand sheet in the United States is identified as the Eureka Quartzite, although its correlatives include the Swan Peak Quartzite and Kinnikinic Quartzite among other formations (e.g., Webb, 1958; Ketner, 1966, 1968, 1977; Poole et al., 1992).
This article is about the petrography, petrology, and diagenesis of the Eureka Quartzite lithosome in the eastern Great Basin in Nevada and Utah (Fig. 1). The Eureka is exposed in erosionally breached folds and thrust sheets that formed during compressional deformation of the Antler and Sevier orogenies. In addition to questions specific to the interpretation of the Eureka are the usual questions about the origin of quartzarenites in general: how were the textural and compositional maturities of such unusual sandstones obtained?; what source areas provided the huge volume of sand?; what was the environment of deposition?; and what was the silica source of all the quartz cement? The Eureka is notable also because it has unusually heterogeneous packing and heterogeneous abundance of quartz cement, and it is the only sandstone of any significant thickness in the study area within a span of 200 million years of deposition.
Vitreous quartzite, quartzite, quartzarenite, and quartz sandstone have all been used to describe beds in the Eureka lithosome (e.g., Kirk, 1933; Webb, 1958; Ketner, 1968; Druschke et al., 2009). Traditionally, the term “quartzite” applies to quartz-cemented sandstones and their metamorphosed equivalents that fracture across—rather than around—detrital quartz grains. Where cementation by quartz was essentially complete (Fig. 2A), Eureka arenite beds are quartzites. Where cementation by quartz was strong but incomplete (Fig. 2B), the beds break around the overgrowths on detrital grains but not across grains. Although locally quartz grains are intensely fractured, there has been no metamorphism. Incomplete cementation in the Eureka yielded many sandstone beds that are weakly friable and porous enough to develop spheroidal weathering (Fig. 3A) and Liesegang bands. In central Nevada, the Eureka is close to and may have been covered by the eastern edge of the Roberts Mountains allochthon (RMa, Fig. 1). The Eureka there is largely quartzite. Only 20 percent of the Eureka in the eastern part of the study area (Utah) is quartzite. I use the term sandstone for all variants. Except for a few feldspar-rich, fine-grained beds that reach subarkose composition and mixed terrigenous–carbonate sandstones, all Eureka samples are quartzarenites (classifications of McBride, 1963; Folk, 1974).
SCOPE OF STUDY
The Eureka Sandstone was examined and sampled at 11 localities in Utah and Nevada (Fig. 1). Samples at some localities were taken systematically through the formation, but at other localities, samples were taken of representative textures and sedimentary structures. Detrital and authigenic minerals and textures were examined using thin (300) and polished (40) sections, binocular microscopy, scanning–electron microscopy (SEM), conventional cathodoluminescence (CL), and scanned CL. Generously, Keith Ketner provided more than 200 thin sections of samples from the Eureka that he acquired during a decade of work (Ketner, 1966, 1968). His samples are evenly spaced from stratigraphic intervals from 14 localities in Nevada and Utah; they include all of my localities.
Because “dust lines” on detrital grains are absent or incomplete on most samples, point counts of thin sections under the petrographic microscope do not accurately assess the volume of quartz cement versus detrital quartz (Evans et al., 1994; Cooper et al., 2000). Thus, point counts were made of 28 samples imaged by scanned CL (Table 1). A grid was superposed on images until 500 points were counted manually per sample. Areas relatively free of through-going fractures were chosen for these data. Several point counts were made of thin sections under the petrographic microscope to determine feldspar abundance (200 counts/sample) in feldspar-rich samples.
δ18O values for both detrital and authigenic quartz and δ18O and δ13C values of dolomite cement are available from a previous study (McBride, 2012). Measurements of porosity using helium and permeability to air were made on 1-inch diameter plugs cut from three samples displaying different degrees of cementation. Permeability was measured under ambient pressure and a confining pressure of 800 psi (Table 2).
In the study area, the sandstone lithosome is known chiefly as the Eureka Quartzite (Fig. 4). In the western part of Utah and eastern part of Nevada, however, the lithosome is divided by a dolostone and shale lens (Crystal Peak Dolostone, up to 50 m thick) into the Watson Ranch Quartzite below (up to 76 m thick) and Eureka Quartzite above (140–180 m thick). The Watson Ranch–Crystal Peak–Eureka triplet reaches a maximum thickness of 200 m in the Ibex (IB) and Crystal Peak (XP) localities in western Utah (Webb, 1958; Hintze and Davis, 2003). The Watson Ranch differs lithically from the Eureka sensu strictu in that it has centimeter- to meter-thick shale beds, greater amounts of detrital illite in the sandstone, more visible bedding, and more abundant trace fossils (cf., Droser and Bottjer, 1989). Both sandstone units within the Eureka lithosome are included in the name “Eureka” in this article. The Eureka lithosome is enclosed in shallow marine dolostone, limestone, and shale (Webb, 1958; Poole et al., 1992; Hintze and Davis, 2003). Shale units several meters thick occur near the top and bottom of the formation at Spor Mountain and in the lower and upper thirds of the formation near Cortez. The Crystal Peak Dolostone is gradational with its enclosing units.
The age of the Eureka in the study area is determined largely from the age of bounding limestone formations, although corals, brachiopods, and conodonts in the Eureka and Crystal Peak Dolostone and carbon isotope stratigraphy confirm its Ordovician age (Staatz and Carr, 1964; Gilluly and Masursky, 1965; Poole et al., 1992; Zimmerman and Cooper, 1998; Hintze and Davis, 2003; Saltzman et al., 2003). Fossils in the bracketing limestone formations indicate that the Eureka includes the late Whiterockia to early Mohawkian stages and that it is younger toward the west (Ketner, 1968; Poole et al., 1992), reflecting regression in that direction. The Whiterockian and Mohawkian stages were originally placed in the Middle Ordovician, the original stratigraphic assignment of the Eureka Quartzite (Ross, 1964). Since 2004, however, the boundary between the Middle and Upper Ordovician series has been relocated to the upper part of the Whiterockian (Ogg et al., 2008), so the Eureka spans the Middle and Upper Ordovician. Deposition possibly took place over a period of no more than 10 million years from 463–453 Ma.
Enclosed within carbonates and shales, the Eureka forms a classical regressive–transgressive clastic tongue whose time duration and geometry categorize it as the clastic part of a third-order stratigraphic sequence (Miall, 2010). Zimmerman and Cooper (1998) and Keller et al. (2003) said that the Eureka in the southern Great Basin is bounded by sequence boundaries and that the regional unconformity at the top of the Eureka is the top of the Sauk IV stratigraphic sequence. The nature of the bounding surfaces of the Eureka, however, is not uniform. Most, but not all stratigraphers (e.g., Saltzman et al., 2003), consider the lower contact of the Eureka with dolostone formations in the study area to be conformable, whereas the upper contact has generally been reported as unconformable (e.g., Poole et al., 1992; Zimmerman and Cooper, 1998; Hintze and Davis, 2003). Druschke et al. (2009) report that the upper contact is conformable at Arrow Canyon (AC); I found it gradational at Cortez (CTZ) and Spor Mountain (SM).
Where the Crystal Peak Dolostone is present, the Eureka is the product of at least two regressive–transgressive cycles. There is a disconformity at the top of an eolian unit within the Eureka at the Sunnyside (SS) and Ibex (IB) localities, and Druschke et al. (2009) recognized an exposure surface in the Eureka in southern Nevada. These hiatal surfaces indicate that the Eureka lithosome does not everywhere record just one or two stratigraphic sequences. Runkel et al. (1998) noted that subdividing Lower Paleozoic quartzarenite formations into stratigraphic sequences is difficult because of the limited range of grain sizes present, because lags of pebbles and granules are needed to identify transgressive surfaces, and because of the scarcity of recognizable features that indicate subaerial exposure. Several geologists have recognized subdivisions of the Eureka based on facies, degree of cementation (sandstone versus quartzite), color (white versus brown), and weathering patterns (Webb, 1958; Ketner, 1968; Miller, 1977; Horn, 1977; Druschke et al., 2009).
BURIAL, THERMAL, AND STRUCTURAL HISTORY
Deposition of the Eureka in the study area occurred about 20° north latitude at a time when the equator was at the position of the Peace River Arch in Canada (Scotese and Golonka, 1992). The area of deposition migrated northward to a position slightly north of the Tropic of Cancer by Jurassic time (Dickinson, 1992).
The burial–uplift history of the Eureka in the study area (Fig. 5) has been reconstructed from published stratigraphic records (e.g., Poole et al., 1992; Miller et al., 1992; Hintze and Davis, 2003). The reliability of this reconstruction is uncertain in view of the structural complexity of the area; stratigraphic sections differ considerably in adjacent mountain ranges.
The Eureka underwent almost continuous subsidence throughout the Paleozoic, although at variable rates (avg. = 27 m/million years). Burial was interrupted during the Antler Orogeny (late Devonian–early Mississippian), when the study area encompassed, from west to east, the Antler foreland basin, forebulge, and back-bulge basin (Ingersoll, 2008). The leading (easternmost) edge of the eastward-moving Roberts Mountains allochthon (RMa) reached approximately the position of my westernmost sample localities in Nevada (Miller et al., 1992; Fig. 1, this study). Although the RMa reached a thickness of 3 km (McFarlane, 1997), mapping suggests that only a thin edge of the allochthon may have covered the Eureka that was sampled (McFarlane, 1997; Saucier, 1997). The Eureka reached maximum burial by the end of the Permian. This depth is estimated from stratigraphic data to be 4.5 km for the area of Millard County, Utah (localities XP, IB) and at least 5 km for localities close to the RMa in Nevada.
The uplift history of the Eureka is poorly constrained as is the thickness of the Mesozoic and Tertiary cover. The Paleozoic rocks, Eureka included, were uplifted and transported southeastward to form an imbricate sequence of thrust plates during the Sevier Orogeny (Jurassic to early Eocene; ∼150–50 Ma; Miller et al., 1992; Cowan and Bruhn, 1992; Willis, 2000). Deformation left a major overprint on the Eureka in the form of folds, faults, fault breccias, quartz-cemented fractures, and high-angle stylolites. Deformation features are greater in Nevada than in Utah. The study area subsequently underwent Tertiary (Basin and Range) uplift, extension, and local igneous intrusion. The Eureka reached elevations greater than 1500 m above sea level during the Tertiary, and its sedimentary cover has been removed in many mountain ranges.
The thermal history of the Eureka Sandstone is also poorly constrained. Present geothermal gradients are highly variable in the study area owing to local igneous intrusions; they range from 25°C/km to >500°C/km in Utah (Blackett, 2004) and from 35–42°C for the area in Nevada (south of Currant Creek, CC) for which data are available (Meissner, 1995). A value of 30°C/km was chosen for a representative thermal gradient of the Eureka throughout most of its history. No samples were collected within 30 km of known igneous intrusions. The assumed gradient and reconstructed burial depths indicate that the Eureka reached a temperature of ∼135°C in Utah and ∼150°C in Nevada.
The light-colored, cliff-forming Eureka sandstone contrasts markedly with the enclosing carbonate formations, which are dark-colored (medium dark gray, N4), fetid, and less resistant to weathering (Fig. 3B). Sandstone beds in the Eureka are mostly some hue of white (N8.5 to N9) to very light gray (N8) (Fig. 3A–C). A few are mottled or uniformly weak red (5Y 5/4) to dark red (7.5R 3/8), and decimeter-size red spots occur locally (Fig. 3D–E). In beds with red color, iron oxide coats detrital quartz grains, or it stains patches of illite. Sandstone outcrops commonly contain coatings of desert varnish of variable thickness, which impart dark grayish brown (10YR 4/2), gray (N5), or black (N2) colors depending on the amount of iron oxide, manganese oxide, and microbial organisms. At the Ibex locality, brown-tinted varnished beds comprise the lower half of the formation (sensu strictu) with stratigraphic continuity (Fig. 3B). At Spor Mountain, 80 percent of the outcrop is covered with brown varnish.
Eureka sandstones display quasi-horizontal master bedding planes spaced from 0.5 to 2 m apart that dominate distant views (Fig. 3C). These laterally extensive bedding planes have millimeter-scale shale partings, a few of which are marked by horizontal trace fossils. Most of the shale partings define millimeter-amplitude stylolites (Fig. 3F). Beds identified by discontinuous bedding planes are only centimeters thick. About 80 percent of the beds at most localities are structureless (Fig. 3A), an observation noted also by Klein (1975) and Miller (1977). Trough and tabular cross-beds, herringbone cross-beds, current-ripple laminations, and horizontal laminations are visible in places as noted by previous geologists, and high-angle, meter-thick cross-beds and translatent ripple laminations are present in rare eolian beds recognized during this study (cf. Kocurek and Dott, 1981). Druschke et al. (2009) describe several localities in southern Nevada and eastern California where bedding is unusually well displayed and where they have recognized decimeter-scale domal stromatolites, wavy laminations of microbial origin, and several near-shore facies. Langenheim and Horn (1978) noted probable stromatolites. Rare load features (Horn, 1977; Miller, 1977), ripples (Miller, 1977), and desiccation cracks (Druschke et al., 2009; this study) have been reported.
The near absence of dark grains (i.e., heavy minerals and rock fragments) and organic-stained clay minerals contribute to the scarcity of bedding structures in the Eureka. Miller (1977) inferred that massive (structureless) beds in the Eureka and other Paleozoic quartz arenite resulted from thorough bioturbation by infaunal organisms, which resulted in the destruction of original bedding. This conclusion is supported in the Eureka by the absence of laminations and the mottled or random distribution of coarse and fine grains at the thin-section scale (Fig. 6).
Skolithos and Teichichnus are the chief trace fossils found in the Eureka (Chamberlain, 1979). Skolithos occurs principally in the Watson Ranch Formation and in the upper 20 m of the Eureka lithosome as scattered burrows (ichnofacies 2 of Droser and Bottjer, 1989) but locally forms stacked “piperock” beds (ichnofacies 4) in units 10 m or so thick. In “piperock” beds, vertical sand-filled Skolithos tubes 2 to 5 mm in diameter and up to 60 cm long almost completely overprint bedding (Fig. 7). In some localities, Skolithos-rich beds alternate with Skolithos-poor beds to form cycles ranging from 80 cm to 1.5 m thick. Chamberlain (1979) found also Arenicolites, Diplocraterion, Saerichnites, and Paleophycus in Nevada exposures. Except as noted above, trace fossils are visible only on bedding planes with clay drapes or in beds where clay or fine-grained carbonate was introduced into sandstone beds during bioturbation.
Secondary megascopic sedimentary features include stylolites (Fig. 3F), iron-oxide Liesegang bands, (former) spheroidal and amoeboid calcite-cemented concretions that have been leached of all their calcite and which now appear as pits (pockmarks of Hintze and Davis, 2003, p. 91; Fig. 8), and heterogeneous patterns of quartz cement. Most stylolites are parallel with bedding, extend laterally less than 1 m, and have less than 2 mm of relief. Stylolites along master bedding planes are meters in extent. In beds with detrital clay or carbonate, stylolite spacing ranges from 10 to 50 cm; in clean beds, stylolite spacing is 1–2 m. Rare, silty, very fine-grained sandstone beds, which have a percent or so of authigenic illite flakes, average five stylolites per centimeter. An insoluble residue of illite, K-feldspar, heavy minerals, and authigenic hematite (alteration of pyrite) mark stylolites. Residue layers are less than 0.5 mm thick except those that formed along clay drapes. A few stylolites occur at a high angle to bedding and reflect a tectonic overprint.
Hemispherical cavities (pockmarks) from pea size to 3 cm in diameter occur in many sandstone beds. The pockmarks comprise up to 30 percent of beds, locally are layer selective, and tend to be coated now with desert varnish. Their size, shape, and distribution are typical of calcite-cemented concretions in sandstone described elsewhere (Ozkan, 2001; McBride et al., 2002). The pockmark hollows in the Eureka are interpreted to have formed by the dissolution in outcrop of the calcite cement that once formed concretions. Amoeboid, decimeter-size patches of weakly cemented sandstone similar to the spherical pockmarks occur at a few localities; these represent irregular-shaped, former calcite-cemented concretions.
Variation in the amount of quartz cement is responsible for beds and parts of beds that appear vitreous (most cement, Fig. 9A), semivitreous, or weakly friable (least cement, Fig. 3A). Some differences are regional (vitreous/semivitreous beds are most common in central Nevada); others are bed specific, lamination specific, or irregular in distribution (Figs. 2B, 9A–B). Spherical quartz-cemented concretions 2–6 cm in diameter (Fig. 9C; nodules of Druschke et al., 2009, p. 1280) are additional examples of the latter feature. They formed where quartz cement completely filled pores within the concretions in contrast to incompletely cemented host sandstone. These concretions have no stratigraphic selectivity and are uncommon, yet widespread, in the Eureka.
Eureka sandstones are chiefly fine- and medium-grained, poorly to well-sorted sand (Trask sorting coefficient from 1.1–1.8; Table 1) composed almost entirely of superbly well-rounded monocrystalline quartz grains cemented by quartz (Figs. 9D, 10; cf. Ketner, 1966). Samples are equally divided between poorly, moderately, and well-sorted textures. Subangular quartz grains are common at the Currant Creek (CC) locality but are otherwise confined to very-fine sand and silt fractions. “Dust lines” on detrital grain margins are incomplete in most samples and completely lacking in others. The majority of samples have a random distribution of finer and coarser grains, a texture typical of bioturbation (Fig. 6). Some Skolithos burrows in piperock are detectable in thin-section by clay-lined burrow walls. Many samples, including those with laminations, are bimodal and consist of a medium sand mode and a fine to very-fine sand mode; coarse sand grains are rare. Less commonly, beds are strongly bimodal (coarse silt–very fine sand versus medium sand) in modes typical of reworked eolian interdune deposits (regs, Folk, 1968). Bimodal beds are poorly sorted. Concave (kidney-shaped) quartz grains (cf. Merino and Hoch, 1988; Werner and Merino, 1997) are common (Fig. 10). Crescentic impact fractures, some filled with hematite, are visible on grains extracted from dolomite-cemented beds that have only incipient microquartz overgrowths (Fig. 11). Molds of mollusks, brachiopods, and corals are present in beds close to contacts with enclosing carbonate beds; local quartz-replaced shells are recognized by their distinctive texture.
Oversize patches of quartz cement, which formed where detrital grains were leached to form secondary pores that were later filled by quartz, are present in all samples (Fig. 12). Oversize cement patches are the size of detrital grains and their enclosing overgrowths. When making point counts of CL images, a judgment was made for each oversize cement patch as to what part of the patch was originally the detrital grain versus pore-filling cement. Lost grains range from 0 to 6.2 percent and average 2.5 percent.
Although the Eureka is composed almost entirely of rigid detrital grains (quartz and feldspar), it has undergone severe and heterogeneous compaction (McBride, 2012). Stylolites, sutured quartz grain contacts (Figs. 9D, 10, 13), fractured quartz grains, and low intergranular volumes (IGV = [sum of volumes of cement and primary porosity] = porosity prior to cementation) are manifestations of Eureka compaction. IGV values are highly variable and range from 9.5–25.5 percent; they are lower close to the RMa (14.4 percent) than to the east (20.8 percent) (Table 1). Although depositional IGV values of 40 percent can reduce to 26 percent by simple grain rearrangement, lower IGV values in sandstone with non-ductile grains require fabric changes by fracturing (Fig. 14) or intergranular grain dissolution (Füchtbauer, 1967; Paxton et al., 2002).
Thin sections from sheared zones are up to 20 percent authigenic quartz. Through-going trains of fluid inclusions and quartz with Boehm lamellae are also fairly common in samples close to the RMa. Brecciated grains that demark millimeter-scale deformation bands and broader fault-gouge zones are typical near the RMa but scarce east of it.
Detrital Framework Composition
The regional detrital framework of the Eureka lithosome is 99.5 percent quartz, 0.5 percent K-feldspar, and trace amounts of carbonate allochems and heavy minerals. Dolomitized carbonate allochems are abundant near contacts with adjacent carbonate formations, where they locally equal quartz grains in abundance, and are a common minor component of the Eureka at the Martin Ridge (MR) and Cortez (CTZ) localities. Rare components are clay clasts, muscovite flakes, collophane clasts, fossil scraps, and chert (five grains in 300 thin sections). Grains lost to dissolution prior to the episode of quartz cementation average 2.5 percent of framework grains. Carbonate allochems were probably the dominant grain type dissolved near carbonate beds at the Cortez and Spor localities, whereas K-feldspar was probably the dominant grain type lost from the Eureka elsewhere. Except for a few feldspar-rich, fine-grained beds that reach subarkose composition and mixed terrigenous–carbonate sandstones, all Eureka samples are quartzarenites (classifications of McBride, 1963; Folk, 1974). A few subarkose sandstones lost sufficient feldspar by dissolution or carbonate replacement to become “diagenetic quartzarenites” (term of McBride, 1987).
Chemical analyses of 17 carbonate-free samples range from 98.6 to 99.9 percent SiO2 (Ketner, 1966, his table 1). Typical of most supermature quartzarenites, there is a near absence of polycrystalline quartz (noted also by Ketner, 1966, and Wu, 1984) and a low abundance of grains with highly undulose extinction. As noted earlier, some quartz grains, especially those in patches of pink and red sandstone, have thin coats of hematite (Fig. 3D–E).
K-feldspar grains are chiefly from 0.05 to 0.07 mm in length (maximum of 0.15 mm). Thus, they are most abundant in the very-fine-grained sandstones (cf. Odom, 1975; Odom et al., 1976). Feldspar is present in only 8 percent of the samples, although it reaches 24 percent of framework grains in one very-fine-grained sample. Feldspar is very well rounded and possesses euhedral prismatic overgrowths. Nearly all is untwinned (Fig. 15). Energy dispersive X-ray data indicate an absence of sodium in the feldspars, so there has been no diagenetic albitization. Detrital cores are cloudier than overgrowths owing to the presence of dissolution micropores. Feldspar is absent in all samples that contain authigenic dolomite; feldspar there was presumably lost to replacement by dolomite. Tabulation on the amount of oversize patches of cement (replaced grains in Table 1) suggests that most samples likely lost from 2–4 percent feldspar by replacement.
Rip-up clasts of sand size are rare, local components in the Eureka. Nearly all clasts are dolomitized limestone. Clay rip-up clasts have been deformed during compaction; they are composed of illite flakes and quartz silt pigmented by organic matter. However, unusual clasts (Fig. 16) in laminations in stromatolite-bearing beds at the Arrow Canyon locality display porous networks of hairs (Fig. 17), fettuccini-like strands, leafy frondescent plates, and face-to-edge plates typical of authigenic clays (Güven et al., 1980; Fig. 18). The clay clasts at Arrow Canyon comprise 15 percent of some beds. These clasts lack detrital quartz grains of silt size and organic matter that are typical of mudstone/shale rip-up clasts but which are typical of bentonite clasts.
Collophane occurs both as well-rounded grains, some with probable organic texture (Ketner, 1966; this study) and as tabular fragments of probable fish remains. The rounded grains have small apatite overgrowths.
Gilluly and Gates (1965) and Gilluly and Masursky (1965) reported the presence of magnetite, staurolite, tourmaline, kyanite, apatite, zircon, brown hornblende, pyroxene, and biotite as detrital heavy minerals in the Eureka, but Staatz and Carr (1964) and I recognized only well-rounded tourmaline (some with overgrowths; Fig. 19) and zircon.
The abundance of detrital clay matrix in Eureka sandstone is problematic because, aside from clay in bioturbated beds and stromatolites, all clay has the texture of authigenic illite or of an illitized precursor clay. Illite flakes that coat detrital grain surfaces, and which are beneath quartz cement, are nearly nonexistent except as micron-size flakes in “dust lines.” Illite whose distribution is typical of a detrital origin occurs mainly in piperock beds and stromatolites (Druschke et al., 2009). Detrital clay in piperock beds, which nowhere exceeds 2 percent, lines the walls of burrows and is dark gray from admixed organic matter that contrasts strongly with the colorless appearance of authigenic illite. Illite is rare on intergranular pressure dissolution surfaces. Characteristics of illite in the Eureka are treated in this paper's section on diagenesis.
Three genetic types of pores are recognized in Eureka sandstones in addition to micropores between clays. In order of decreasing abundance, these are narrow, linear circumcement pores, dissolution pores, and normal primary mesopores. [The size terms devised by Hendry and Trewin (1995) for authigenic quartz crystals are here applied also to pore widths measured in SEM images: macropore = >50 μm, mesopore = 10–50 μm, micropore = <10 μm.] Although circumgranular pores up to a micron wide separate one crystal from another in nearly all rocks (including granite) imaged with the SEM, these pores in other than vitreous Eureka samples are typically much wider (Fig. 20). In beds with meso and macrocrystalline cement, overgrowths are separated from adjacent overgrowths by linear circumcement pores. In vitreous quartzite, the best indurated of Eureka sandstone beds, these pores are less than 2 μm wide (Fig. 21). In friable Eureka samples, these pores range from 5–20 μm wide (Fig. 20). Quartz exposed in the walls of circumcement pores is smooth and unetched, indicating that the pores are the product of incomplete cementation and not dissolution of cement in the outcrop. Thus, they are primary pores. Circumcement pores do not exceed 2 percent of rock volume (Table 2). Circumcement mesopores are more abundant in Utah than in Nevada.
Porosity and permeability measurements made on 1-inch diameter plugs of semifriable, semivitreous, and vitreous samples show that measured porosity does not exceed 1.9 percent and permeability to air does not exceed 0.042 md (Table 2). The semifriable sample has four times greater porosity than the vitreous sample and two orders of magnitude greater permeability. Porosity measured under stress of 800 psi is 10 percent lower than for unstressed samples.
Near-surface leaching of dolomite cement and dolomitized allochems have generated up to 8 percent dissolution porosity. Other than circumgranular pores, primary mesopores comprise less than 2 percent of samples except for those that have a strongly heterogeneous distribution of quartz cement. Rare, poorly cemented laminations reach 22 percent porosity (Fig. 2B).
Authigenic phases other than quartz and dolomite occur in trace or minor amounts. These include K-feldspar, illite, pyrite, iron oxide, kaolinite, tourmaline, apatite, opal, and pyrobitumen (“dead oil”) (e.g., Gilluly and Masursky, 1965; Ketner, 1966, 1968; Wu, 1984; this study). In addition, several features point to the previous existence of sparse “early” calcite cement. These include: 1) pockmarks (Fig. 8), in whose margins intergranular volume (IGV) values exceed 30 percent; 2) ghosts of calcite and aragonite allochems that developed carbonate overgrowths, all of which are now composed of sparry dolomite; and 3) concentration of dolomite cement at the margins of dolostone beds where calcite and aragonite cementation would first occur. The relatively high IGV values of pockmark margins attest to an episode of carbonate cementation prior to the main phase of quartz cementation.
Illite pseudomorphs of probable authigenic kaolinite (Fig. 22) occur at Arrow Canyon (AC) and Ibex (IB). Contact metamorphic minerals adjacent to igneous intrusions near Cortez include calcite (dedolomite), talc (Gilluly and Masursky, 1965), epidote, clinopyroxene, amphibole, and kaolinite (this study). Calcite and dolomite also occur as trace amounts in caliche; iron oxide and manganese oxide occur in desert varnish.
Quartz cement dominates throughout the Eureka except near contacts with dolostone beds, where dolomite, which inferentially replaced calcite and aragonite allochems and filled adjacent pores, is dominant. Generally, quartz is the only cement present except for traces of illite. Cement volume throughout the study area ranges from 9.5–25.5 percent (Table 1). It is generally lower in central Nevada (mean = 14.4 percent) than farther east (mean = 20.8 percent); its abundance is largely controlled by the pore space that survived compaction (McBride, 2012).
Heterogeneous cementation by quartz is widespread at the outcrop-, hand specimen-, and thin-section scale (Figs. 2B, 9B). Two degrees of cementation (friable, semivitreous, or vitreous samples) are commonly present in the same outcrop, although one style is dominant. Friable samples have a milky appearance (Fig. 3A), vitreous samples appear glassy (Fig. 9A, B), while semivitreous samples are intermediate in appearance. Where semivitreous cement is dominant, vitreous cement occurs as stratiform lenses (Fig. 9A), spheroidal concretions (Fig. 9C), or irregular patches (Fig. 9B). A few beds display strongly selective cementation of finer-grained laminae (Fig. 2B). In one bed with translatent ripple bedding, coarse-grained laminations retain 22 percent porosity, but fine-grained laminations retain only 2 percent porosity. Only a crust of microcrystalline quartz overgrowths cements coarse laminae in this sample, whereas fine laminae are cemented by both microquartz and mesoquartz. These gross heterogeneities are at the decimeter scale. For many samples with variable amounts of cement, there is no obvious textural or compositional control of the cement. An exception is the presence and abundance of illite (see below). Outcrops do not show any correlation between position of stylolites, which are potential sources of silica for quartz cement, and megascopic degree of cementation.
Quartz cement is a combination of microquartz and mesoquartz (80 percent) and minor macroquartz (20 percent) (Fig. 23). Typical of quartz-cemented sandstones, cementation in the Eureka began as microcrystals (microquartz) that grew into mesocrystals and, in places, into macrocrystals (cf. Waugh, 1970; Pittman, 1972). In the Eureka, however, growth of microquartz and mesoquartz persisted much longer than normal, and only in the last stage of cementation did macroquartz form in some samples (Fig. 23). Some laminations did not progress beyond the microquartz/mesoquartz stage (Figs. 24, 25), and microquartz is the only quartz cement in some pockmarks. Flakes of illite are ubiquitous in samples with exposed microquartz overgrowths. Illite coats some microcrystal faces and is intergrown with a variety of orientations with others (Fig. 26). Macroquartz is the most abundant cement only in samples where illite has exceptionally sparse occurrence.
Cathodoluminescence (CL) imaging rarely distinguishes macroquartz from microquartz. Some microquartz luminesces blue and mesoquartz luminesces red. Sector zoning (Fig. 12) is independent of cement texture. Microquartz and mesoquartz are recognizable only in thin sections where they are the only cement on a grain.
Rare tiny (<1 μm) euhedral quartz crystals occur both as “outgrowths” on the surface of mesoquartz facets and as free-standing crystals in pores; both are of late-diagenetic origin. Chalcedony cement occurs in a few vitreous beds. Microquartz grains in chalcedony are smaller than 0.5 μm.
Fractured quartz grains and grains in deformation bands and fault breccia at the thin-section scale have been healed by non-luminescing (black and white CL) and dark blue (color CL) quartz cement. Through-going fractures transect quartz cement (Fig. 27).
Other Authigenic Phases
Cloudy dolomite spar crystals replaced mollusks or brachiopods and other unidentifiable carbonate allochems and their overgrowths and also occur as pore-filling cement in allochem-free patches and beds. The dolomite has the coarse euhedral/subhedral morphology and distribution of typical replacement dolomite. Dolomite etched microquartz overgrowths. Some dolomite is weakly zoned under CL. δ18O values range from −5 to −7.6‰ and δ13C values from 0.7 to 5.7‰. Isotope values are similar to those for the Crystal Peak Dolostone (McBride, 2012). The fetid odor of dolostone beds indicates that the dark color of dolomite grains is imparted by hydrocarbon and H2S inclusions. Some dolomite is weakly zoned under CL.
K-feldspar overgrowths comprise a trivial volume of Eureka samples (Fig. 15). Overgrowths embay microquartz cement and, thus, were partly simultaneous with microquartz.
Authigenic illite is visible in all samples examined with the SEM, although it cannot be seen in all thin sections, especially those of vitreous samples. Single flakes (less than 4 μm long and 0.05 μm thick) tend to be perpendicular to detrital grain surfaces. Illite flakes in clusters have a random orientation and are also intergrown with and coat microquartz crystals (Fig. 26). Clusters reach 10 μm thick and 20 μm long; flakes tend to be well aligned.
Clay in most Arrow Canyon samples has the appearance in thin section of colorless detrital matrix or of clay clasts that have been mashed to fill and conform to the shape of intergranular pores (pseudomatrix, Fig. 16). SEM images show, however, that illite in these samples is present as chaotically arranged flakes with 60-percent porosity within the clay, there is an absence of detrital quartz silt and organic matter, and illite is intergrown with microquartz. In addition to typical flakes, illite occurs in the unusual textures described earlier (spiky ends, laths, hairy mats, and vermicular rosettes; Figs. 17, 18, 22). The former three are textures of authigenic illite (Güven et al., 1980; Wilkinson and Haszeldine, 2001) and the latter of illitized kaolinite. Although most detrital quartz surfaces have well-developed microquartz crystal coats, some surfaces in contact with illite have only a minute layer of flat quartz platelets. The textures suggest that illite in these samples formed either as a direct pore fill or by recrystallization of clay clasts/pseudomatrix. Where clay clasts were in close contact with detrital quartz grains, or before the clay clasts were mashed, only a trace of overgrowth was able to form. Away from clay clast–quartz grain contacts, typical microquartz overgrowths formed. The vermicular illite has the morphology of authigenic kaolinite and appears to be a pseudomorph of it.
Gilluly and Masursky (1965) and Ketner (1966, p. C56) reported pyrite cubes, but I found only their hematite pseudomorphs. The latter are present in approximately 15 percent of the thin sections, where they range from a trace to 2 percent. Cubic pseudomorphs and hexagonal crystals of hematite reach 0.2 mm, and crystal clusters reach 0.5 mm. Red hematite occurs as micron-sized dust particles on detrital quartz-grains beneath quartz overgrowths (Fig. 3E), within fractures in detrital quartz, as unresolved specks in the insoluble residue of stylolites, in millimeter-scale haloes around pyrite pseudomorphs, and as grain coatings in Liesegang bands. SEM imaging shows that the pyrite crystals embay both detrital and authigenic quartz. Brown iron oxide, in addition to manganese oxide, is widespread as desert varnish. Of 17 samples of dolomite-free samples of Eureka from Sunnyside, Nevada, Ketner (1966, his Table 1) found Fe2O3 content to range from 0 to 0.22 percent and FeO to range from 0.4 to 1.0 percent.
Some tourmaline grains developed ragged overgrowths up to a third the length of their detrital core (Fig. 19). Such overgrowths on tourmaline are typical of supermature quartzarenites (the author's sandstone collection). Thin crusts of opal fill the capillary pores in a few samples in Utah. The opal is likely the product of Tertiary or Quaternary weathering, where silica was derived from microquartz in the Eureka. Phosphate (collophane) grains have minute euhedral crystal outgrowths. “Dead oil” (pyrobitumen) that coats microquartz crystals occurs in the Eureka at the Sunnyside (SS) and Crystal Peak (XP) localities.
CRYSTAL PEAK DOLOSTONE
Dolostone beds in the Crystal Peak Dolostone resemble those that enclose the Eureka lithosome. They are dark-colored (medium dark gray, N4), fetid, partly chertified (stringers and nodules), and riddled with stylolites. Beds are less than one-half meter thick, and they are separated by centimeter-scale shales. Most beds are burrow-mottled or structureless, although those close to sandstone beds of the Eureka lithosome have ripple laminations and trough cross-beds.
Except for rippled grainstone beds, dolostone beds are either fossiliferous mudstones or packstones, and many contain several percent well-rounded quartz sand grains typical of Eureka samples. Many fossils (bryozoans, brachiopods, trilobites, gastropods, algae) have been replaced by sparry dolomite, whereas original carbonate mud is now dolomicrite or dolomicrosparite. Dolomitization obscures much of the original texture of the rocks, and only locally did a few sparry calcite crystals survive replacement.
Carbon and oxygen isotopic values for dolostone in the Crystal Peak and the dolostone underlying the Eureka at the Whipple locality have values similar to those for cement in the Eureka. δ13C ranges from −0.7 to −5.6 (PDB) and δ18O from −6.0 to −7.6 (VSMOW) (McBride, 2012).
The Eureka displays many characteristics of other Early Paleozoic quartzarenites in terms of its geometry, mineralogical and textural maturity, and sedimentary features (e.g., Dapples, 1955; Heald, 1956; Ketner, 1968; James and Oaks, 1977; Barnes et al., 1992; Dott, 2003, Runkel et al., 2007). It is distinctive in its heterogeneous abundance of quartz cement, wide range of anomalously low IGV values, evidence of feldspar replacement, presence of former calcite-cemented concretions, and relicts of former red sandstone beds.
Various source areas have been suggested for the sand in the Eureka (summaries in Ketner, 1968; Poole et al., 1992; Gehrels and Dickinson, 1995). From grain-size and sorting trends and paleogeologic reconstructions, Ketner (1968) interpreted the Eureka and sister formations to be composed of quartz grains reworked from Cambrian sandstones situated on the Peace River–Athabasca Arch in Canada. Sand-rich middle Cambrian rocks more than 100-m thick once covered the arch (van Hees, 1964). U–Pb dating of zircon grains in the Eureka from near Ely, Nevada, indicates that basement rocks of the Peace River–Athabasca Arch area were the ultimate source of detritus for those samples (Gehrels and Dickinson, 1995). The absence of quartz grains with abraded overgrowths in the Eureka indicates either that pre-Ordovician sandstones that provided detritus to the Eureka had not undergone cementation by quartz or the unlikely event that once-present overgrowths were removed from detrital grains by abrasion. Although recycling of older sandstones is the most common origin of quartzarenite (Dott, 2003), the Eureka lacks such evidence. Quartz grains with superb rounding in the Eureka are indicative of an intense eolian abrasion history. Such grains may have been derived from Proterozoic sand-seas that existed north of the Peace River–Athabasca Arch in the Northwest Territories (Ross and Donaldson, 1982).
Zircon ages of the Kinnikinic Quartzite in Idaho, correlative in part to the Eureka, indicate contributions of detritus from the Transcontinental Arch and other local sources (Pope et al., 2008). Given the volume of sand in the Eureka, multiple source areas were likely.
The scarcity and even absence of polycrystalline and undulose quartz is a feature of supermature quartzarenites in spite of their ultimate provenance from granitic-rich igneous and metamorphic terranes. Polycrystalline quartz has great durability (Harrell and Blatt, 1978) but undergoes disaggregation to silt-size particles during chemical weathering (McBride and Abdel-Wahab, 1996). Undulose grains are inferentially less durable than non-undulose grains, but experimental evidence is lacking.
Sediment Transport History
Sand grains in the Eureka were likely transported to their depositional site by a combination of coastal, eolian, and fluvial processes. However, the combination of textural features (such as the superb rounding of quartz coarser than silt size, abundance of bean-shaped grains, bimodal grain sizes typical of reworked eolian regs, and crescentic impact scars) indicates that most sand grains that make up the Eureka had a long eolian transport and abrasion history. In addition, the local iron-oxide coats on quartz grains in the Eureka are typical of sands in mature eolian ergs (Folk, 1976; Walker, 1979; and references therein) and are compatible with marine Eureka sand having been derived from such erg deposits. During Whiterockian and Mohawkian time, the Eureka depositional site was below 30 degrees north latitude and within the arid climatic belt for some time (Scotese and Golonka, 1992); thus, the opportunity for the development of red erg sands existed.
One wonders how much of the Eureka might have been red at the time of deposition. The amount of pyrite (locally 2 percent, now all hematite) in the Eureka in the absence of organic matter is anomalously high for clean sandstone deposited in shallow, well-oxygenated seawater. Much of the iron in pyrite in the formation probably was derived from the reaction of hematite grain coats with hydrocarbons and H2S that invaded the formation at depth. Ferric iron, reduced in the presence of such fluids, will react with H2S to form pyrite (cf. Levandowski et al., 1973; Surdam et al., 1993; Beitler et al., 2003; Parry et al., 2004). The presence of hematite within fractured quartz grains (including crescentic eolian impact fractures) in otherwise hematite-free beds is compatible with the hypothesis that such isolated hematite survived the reduction event that chemically removed once-extensive hematite grain coats. Scattered red spots in otherwise white beds are other likely survivors of such a reduction event. How extensive such a bleaching event might have been is unknown.
Kuenen (1959, 1960) argued, without refutation so far, that his experimental abrasion studies indicated that well-rounded quartz grains of sand size could only be generated in abundance by abrasion in air and not by abrasion in the denser medium of water. Eolian abrasion is apparently essential to round quartz grains of medium and finer sand size. The preservation of crescentic impact scars of eolian origin on quartz grains shows that abrasion in a marine environment of deposition (see below) was not strong enough to obliterate all such eolian features. The same conclusion applies to the coating of iron oxide present on some quartz grains. Although such coatings, inherited from desert dunes (Folk, 1976; Walker, 1979; satellite images), tend to be removed by abrasion (Gardner and Pye, 1981), some survive to burial (Ozkan, 2001; McBride et al., 2002).
The survival of K-feldspar, albeit chiefly in the fine–very fine sand fraction, attests to its mechanical durability. The ability of K-feldspar to avoid albitization at maximum burial depths indicates that only chemically stable grains survived weathering of its source rock and subsequent abrasion.
Environment of Deposition
The abundance of well- and moderately sorted sand, presence of herringbone and current-rippled cross-beds, abundance of shallow-marine fauna and trace fossils, and scarcity of shale led previous geologists to interpret the Eureka to be shallow marine (subtidal, intertidal, and shoreface) in origin (Gilluly and Gates, 1965; Ketner, 1966, 1968; Klein, 1975; Horn, 1977; Miller, 1977; James and Oaks, 1977; Deen, 1984; Wu, 1984; Druschke et al., 2009). Piperock beds, whose burrows were likely made by polychaetes and phoronids, are also chiefly littoral, beach, and intertidal deposits (Droser, 1991). The absence of hummocky cross-beds indicates that water depths did not exceed storm wave base. Gilluly and Masursky (1965, p. 17) mentioned possible subaerial deposits of unspecified origin near Cortez, Nevada. Wu (1984) mentioned possible eolian deposits in central Nevada. I recognize a 20-m-thick sequence of meter-scale eolian cross-beds with translatent ripple laminations in the upper third of the formation at the Sunnyside (SS) and Ibex (IB) localities (Fig. 1). However, typical of Early Paleozoic supermature quartzarenite that had a long eolian history, the wind-abraded quartz grains of the Eureka ended up largely in shallow-marine environments (Grabau, 1913; Twenhofel and Thwaites, 1919; Dake, 1921; Ozkan, 2001; McBride et al., 2002; Dott, 2003). The considerable thickness of the Eureka and its younging toward the southwest indicates that it is not simply the product of flooding of an erg but is a regressive deposit fed from a huge supply of sand. Druschke et al. (2009) provide the most detailed coastal environmental reconstruction of the Eureka exposed in southern Nevada and eastern California.
The poor to moderate sorting of many sandstone beds is not typical of deposits of so-called “high energy environments.” However, bioturbation can not only destroy primary laminations, but it can also mix well-sorted coarser and finer layers into poorly sorted mixtures.
Cause of Eureka Regression
Deposition of the huge volume of Eureka sand during Whiterockian and Mohawkian time was a major anomaly in sedimentation on the Cordilleran miogeocline from Late Cambrian through Late Devonian because, except for the Eureka, this time span records the accumulation of essentially only carbonate and minor shale. The report of Middle to Late Ordovician glaciation in Gondwana of 35 million years duration (Frakes et al., 1992) raised the possibility that the regressive Eureka was deposited in response to the glacioeustatic lowering of sea level. Subsequent to the study of Frakes et al. (1992), dispute arose over the age and duration of Gondwana glaciation and whether the several Late Ordovician regressive sequences on the North American craton are synchronous (e.g., Brenchley et al., 1994; Sutcliffe et al., 2000; Pope and Steffen, 2003; Mitchell et al., 2004). Until these questions are resolved, whether the Eureka regression was the product of glacioeustasy, non-glacial eustasy, or epirogenic uplift of the North American craton remains unanswered. The eolian deposits interbedded with marine deposits in the Eureka, shown also by other Lower Paleozoic quartzarenites (Nadon et al., 2000), identify small-scale stratigraphic sequences, but these require sea-level changes of only meters.
Table 3 lists the diagenetic events recognized for the Eureka and their inferred timing. Figure 5 shows the more important diagenetic events tied to the burial and uplift history of the Eureka. The scenario below is for the Eureka in Utah. The Eureka was buried deeper in central Nevada, where it has more completely cemented beds and is more fractured.
Shallow Burial: Ordovician–Silurian: Burial to ∼1 km
Compaction and precipitation of small amounts of authigenic phases characterize the shallow burial phase of the Eureka. At such shallow depth and in the near absence of ductilely deformed grains, compaction was almost entirely by grain rearrangement. However, intergranular pressure dissolution of silicates can occur as shallow as 900 m (McBride et al., 1991), and fragile carbonate fossils can be crushed. Initial porosities of 40–45 percent were likely reduced to ∼30 percent by burial to 1 km (cf. McBride et al., 1991; Gluyas and Cade, 1997; Paxton et al., 2002).
The early diagenetic phases shown in Table 3 have too few textural clues to determine their relative ages. The abundance and large size of euhedral K-feldspar overgrowths indicate that they formed at shallow depth before compaction significantly reduced pore size. Components for K-feldspar overgrowths and authigenic kaolinite probably came from the dissolution of the unstable K-feldspar grains (cf. Thein, 1966; Stablein and Dapples, 1977; Bjørkum and Gjelsvik, 1988; McBride et al., 2002). There are no textural clues to the source of components for overgrowths of collophane and tourmaline; they probably also were derived from unstable sister grains. Ketner (1966) noted that some pyrite may have been syndepositional. The hypothesis proposed here, however, is that most pyrite formed during invasion of hydrocarbon and H2S at considerable depth.
Medial Burial: Devonian: Burial from ∼1km to ∼2.5 km
This phase of diagenesis was dominated by strong compaction, initial precipitation of microquartz and illite, and patchy cementation by carbonate, probably calcite. Calcite components were likely derived from the dissolution of aragonite biogenic grains in nearby carbonate sediments. By burial depths of 2–2.5 km, quartzose sands generally reach IGV values of about 26 percent, at which point strong, stable sand frameworks have formed (cf. Füchtbauer, 1967; McBride et al., 1991; Stone and Siever, 1996; Paxton et al., 2002). At an IGV of 26 percent, initial porosities in the Eureka would have been reduced by 35–44 percent.
Compaction of the Eureka from overburden was probably enhanced by compressional stresses of the Antler Orogeny and, in central Nevada, possibly by the weight of the eastern edge of the RMa.
IGV values below 26 percent in sandstones lacking ductilely deformed grains require chemical compaction (pressure dissolution) and/or fracturing of brittle grains to attain such values (cf., Dunn, 1993; Onasch, 1994; Paxton et al., 2002; Makowitz and Milliken, 2003). The greater abundance of fractured grains and lower IGV values in central Nevada than in samples farther east probably result from the contribution of grain fracturing to the compaction process.
The data sets of McBride et al. (1991), Gluyas and Cade (1997), and Paxton (2002) indicate that depths of 2.5 km in samples can attain IGV values slightly below 26 percent where rigid grain compaction is dominant. Stone and Siever (1996), however, report an IGV value of 17 percent at only 1.5-km burial. In contrast with the Eureka, many of their samples contain clay coats, which are known to enhance grain dissolution (Heald, 1959; de Boer et al., 1977; Bjørkum, 1996). The approximately 180 million years it took the Eureka to reach 2.5 km of burial was sufficient to generate the low IGV values it attained. The greater compaction shown by Eureka samples in central Nevada compared to those farther east reflects closer proximity of the former localities to the Antler and Sevier forelands and possibly burial beneath the RMa and Sevier thrust sheets.
The wide range of IGV values found at single localities (e.g., SS in Table 1) indicates that compaction proceeded unevenly and that hydrostatic pressure compartments (i.e., Ortoleva, 1994) were not important. Two features that could influence differential compaction are patchy cementation, where early cement would retard compaction, and patchy illite coatings on quartz grains, where illite would locally enhance intergranular dissolution. Coats of illite on quartz grains are essentially absent except for Skolithos-rich piperock beds and stromatolites at Arrow Canyon (cf. Druschke et al., 2009), so this explanation is rejected. If the initial cementation by quartz was of variable abundance and led to differential compaction, as seems essential, the evidence has been buried by later cementation. Although locally finer-grained laminations have more cement than coarser laminations, laminations are not abundant in the Eureka. The most likely reason for the disparity in IGV values is uneven cementation by quartz, which, in turn, was controlled by the abundance of illite cement (McBride, 2012). If trace amounts of detrital illite were shuffled around unevenly during bioturbation, they acted as seed crystals for illite cement, but the evidence is also buried beneath later quartz cement. The main episode of quartz and illite cementation took place during deep burial, but the onset of precipitation of both phases was likely during intermediate burial.
Dissolution pores, now filled by quartz cement, formed by the dissolution of feldspar and possibly skeletal carbonate grains. Dissolution was at close to maximum burial depth, or the dissolution pores would have collapsed by compaction.
Deep Burial: Mississippian–Cretaceous: Burial from ∼2.5 km to 4.5 km
Deep burial diagenesis was dominated by simultaneous precipitation of quartz and illite cement and the development of most strain features (Boehm lamellae, deformation bands, fault gouge) during the Antler and Sevier orogenies. Through-going fractures that cut quartz overgrowths—but are now cemented—show that some quartz grains were fractured and healed after the main episode of cementation. The temperatures at depths greater than 2.5 km (>95°C) was high enough to speed up quartz cementation (Giles et al., 2000; Walderhaug, 1996; Walderhaug et al., 2000; Worden and Morad, 2000), and this, in turn, would retard and eventually halt compaction. The δ18O values obtained for quartz cement (+16.5–22‰) do not permit calculating a unique composition of the cementing fluid. Values are compatible with either a residual evaporative brine or meteoric water modified by diagenetic reactions. Evaporites, now lost to dissolution, may have been deposited in the arid setting of Eureka dune sands; or meteoric water, recharged from the tectonic highlands to the west, may have reached the Eureka during deep burial.
Illite ceased growth prior to the phase of macroquartz cementation. The components for growth of illite were probably derived from renewed dissolution of unstable K-feldspars in sandstone and shale beds. The clast-size patches of pure illite at Arrow Canyon differ from clay rip-up clasts and shale clasts in the absence of either iron oxide or organic matter. These clasts originally were probably pure clay, either smectite (bentonite) or kaolinite clasts that were altered to illite. The rare vermicular illite grains are probably replaced authigenic vermicular kaolinite that formed during shallow-burial diagenesis.
The precipitation of authigenic illite and prolonged precipitation of microquartz served to retard and, in places, prevent the development of normal meso and macroquartz overgrowths on the Eureka by at least two processes. First, microquartz cementation proceeds slower than meso and macroquartz (Heald and Renton, 1966), commonly by more than an order of magnitude (Bonnell et al., 2006; Lander et al., 2006; Lander et al., 2008). Secondly, illite that coated microquartz crystals prevented further quartz precipitation at those sites (McBride, 2012).
When illite growth stopped, probably due to the depletion of unstable K-feldspars, meso and macroquartz precipitation proceeded in most beds (Fig. 23). Growth of macroquartz reached different stages, but it was favored in the western part of the study area where burial and temperature were greatest. The failure of cementation to go to completion in many samples left the distinctive circumgranular pores common in Utah samples. If the burial–thermal histories are correct, the failure of cementation to reach completion must reflect a shortfall of silica in solution. The scarcity of illite in the strongest-cemented beds and its abundance in the weakest-cemented beds suggest that the abundance of illite was the chief control of quartz cement abundance. Illite is sparse in beds with the least feldspar, which suggests that original feldspar abundance was the primary influence on illite abundance. The development of spheroidal, tabular, and irregular-shaped quartz-cemented concretions is unusual and known elsewhere only in silcrete (Thiry and Maréchal, 2001; K. Milliken, personal communication, 2003; my data). Their controls are unknown.
Commonly cited sources of silica for quartz cement in sandstones, such as feldspar dissolution, smectite-to-illite conversion in thick underlying shales, pressure dissolution at quartz–quartz grain contacts, and siliceous sponge spicules (McBride, 1989), do not apply to the Eureka. Consequently, the dissolution of quartz at intraformational stylolites was suggested as the most probable silica source in the Eureka (McBride, 2012). Stylolites have the potential to yield large volumes of silica for quartz cement (Dutton and Diggs, 1990; Stone and Siever, 1996). There is no accurate way at present, however, to quantify the amount of silica produced by such dissolution, especially stylolites at sandstone–shale contacts (Stone and Siever, 1996; Spötl et al., 2000). Thus, the stylolite hypothesis for the Eureka has not been tested. It is probably not coincidental, however, that the weakest-cemented samples have the sparsest stylolites.
Oxygen isotopic values of dolomite cement and beds suggest that dolomitization took place at elevated temperatures and was a regional rather than a local process such as reflux. Fluids expelled from the Antler and Sevier orogens were possible dolomitizing fluids. Carbon isotopic values indicate that the carbon was chiefly derived from recycled skeletal grains.
Uplift and Exposure: Cretaceous–Present: Uplift to ∼2 km Above Sea Level
Beginning near the end of the Cretaceous, a long period of uplift and erosion resulted in the exposure of the Eureka in folds and thrust sheets whose current elevations reach nearly 2 km. Close to the surface, invasion by meteoric water permitted: the oxidation of pyrite to hematite; the precipitation of some hematite in Liesegang bands; the dissolution of dolomite in concretions and scattered cement grains; local precipitation of opal cement; and widespread, but small-scale, precipitation of caliche calcite and widespread formation of desert varnish.
Silica for opal was probably generated by dissolution of cement (chiefly mesoquartz) in the Eureka; Ca in caliche and Mn and Fe for desert varnish came from the carbonate beds within and adjacent to the Eureka. Some Fe in desert varnish probably also came from pyrite, iron oxide, or iron-bearing dolomite in Eureka sandstones.
This regional study of the Eureka Quartzite leads to the following conclusions:
1) Deposition of the huge volume of Eureka sand was a major anomaly in sedimentation on the Cordilleran miogeocline from Late Cambrian through the Late Devonian because, except for the Eureka, this time span records the accumulation of essentially only carbonate and minor shale. Whether the westward Eureka regression was the product of eustasy or epirogeny remains unanswered.
2) Quartz and feldspar sand grains were transported from their source terrains as far north as the Peace River–Athabasca Arch in Canada and as far east as the Transcontinental Arch in central North America by several processes, but the shape, roundness, and surface texture of the quartz grains were generated by prolonged eolian abrasion in desert ergs.
3) Most iron-oxide grain coats, formed in eolian dunes, were removed by marine abrasion plus contact with reducing diagenetic fluids. The reduced ferric iron survived in pyrite, which was subsequently oxidized in outcrop to hematite. Judging from the abundance of former pyrite, a significant part of the Eureka sandstone was originally red.
4) The Eureka was deposited in shallow-marine and intertidal environments except for rare, meters-thick eolian sand bodies. Extensive bioturbation destroyed most of the primary stratification and worsened sorting by mixing initially well-sorted beds of different grain size.
5) The Eureka, although composed almost entirely of rigid framework grains, underwent strong compaction as shown by low IGV values. Compaction was more important than quartz cementation in destroying primary porosity, but the combination of processes has reduced reservoir quality to that of a “tight gas sand.”
6) Compaction took place by the combination of grain rearrangement, pressure dissolution, and grain fracturing. The wide range of IGV values within the same outcrop reflects uneven compaction owing to heterogeneous quartz cementation.
7) Quartz cement abundance is variable at scales ranging from the outcrop, hand specimen, and thin-section. Strongly cemented samples are vitreous; incompletely cemented samples are weakly friable.
8) Microcrystalline and mesocrystalline quartz overgrowths comprise 80 percent of the quartz cement in contrast with 20 percent macroquartz. Some samples have only micro and mesoquartz cement. The most likely source of most quartz cement was quartz grain dissolution at intraformational stylolites.
9) Authigenic illite is pervasive, although volumetrically minor. Dissolution products of K-feldspar likely provided components for the illite. Illite flakes coated quartz overgrowths and impeded complete cementation by quartz in places.
10) All detrital grains of K-feldspar have overgrowths. The abundance and size of euhedral overgrowths indicate that they formed at shallow burial depths and before major compaction. Overgrowths of likely shallow-burial origin occur also on collophane and tourmaline.
11) Recrystallized clay clasts at Arrow Canyon that are composed of white authigenic illite unpigmented by either iron-oxide or organic matter were either originally bentonite or kaolinite.
12) Alteration of Eureka sandstones continued after uplift and exposure. Events include the oxidation of pyrite, dissolution of carbonate, precipitation of traces of opal cement, and precipitation of manganese and iron oxide in desert varnish. Dissolution of carbonate-cemented concretions left abundant pockmarks.
The following provided assistance for this study: Antar Abdel-Wahab, Tim Diggs, Shirley Dutton, Tony Ekdale, Robert Folk, Jeffrey Horowitz, Keith Ketner, Lynton Land, Larry Mack, Kitty Milliken, Laura Net, Aysen Ozkan, M. Dane Picard, and Rob Reed. Shirley Dutton and Kitty Milliken provided constructive comments on an earlier version of the manuscript. The present manuscript was measurably improved by comments by reviewers Bob Dott, M. Dane Picard, and RMG co-editor Arthur Snoke.
- Received June 6, 2012.
- Revision received September 1, 2012.
- Accepted September 11, 2012.