The Late Ordovician (Sandbian) Glasford structure: A marine-target impact crater

with a possible connection to the Ordovician meteorite event

Charles C. MONSON 1,2*, Dustin SWEET 3, Branimir SEGVIC 3, Giovanni ZANONI 3,Kyle BALLING2, Jacalyn M. WITTMER2,4, G. Robert GANIS5, and Guo CHENG6

1Illinois State Geological Survey, University of Illinois at Urbana-Champaign, 615 East Peabody Drive, Champaign, Illinois61820, USA

2Department of Geology, University of Illinois at Urbana-Champaign, 1301 West Green Street, Urbana, Illinois 61801, USA3Department of Geosciences, Texas Tech University, 1200 Memorial Circle, Lubbock, Texas 79409, USA

4Department of Geological Sciences, State University of New York at Geneseo, 1 College Circle, Geneseo, New York14454, USA

5Consulting Geologist, 749 Burlwood Drive, Southern Pines, North Carolina 28387, USA6Department of Earth and Environmental Sciences, 115 Trowbridge Hall, Iowa City, Iowa 52242, USA

*Corresponding author. E-mail: cmonson@illinois.edu

(Received 09 August 2018; revision accepted 16 September 2019)

Abstract–The Glasford structure in Illinois (USA) was recognized as a buried impact craterin the early 1960s but has never been reassessed in light of recent advances in planetaryscience. Here, we document shatter cones and previously unknown quartz microdeformationfeatures that support an impact origin for the Glasford structure. We identify the 4 kmwide structure as a complex buried impact crater and describe syn- and postimpact depositsfrom its annular trough. We have informally designated these deposits as the KingstonMines unit (KM). The fossils and sedimentology of the KM indicate a marine depositionalsetting. The various intervals within the KM constitute a succession of breccia, carbonate,sandstone, and shale similar to marine sedimentary successions preserved in other craters.Graptolite specimens retrieved from the KM place the time of deposition at approximately455 � 2 Ma (Late Ordovician, Sandbian). This age determination suggests a possible linkbetween the Glasford impact and the Ordovician meteorite shower, an increase in the rateof terrestrial meteorite impacts attributed to the breakup of the L-chondrite parent body inthe main asteroid belt.

INTRODUCTION

The Glasford structure is a 4 km wide buried impactstructure located in west-central Illinois, USA (Fig. 1). Itwas identified as a probable impact structure in 1963(Buschbach and Ryan 1963; Ryan and Buschbach 1963)but subsequently went largely ignored by the scientificcommunity, apart from identification of shatter cones inthe 1980s (McHone et al. 1986; Sargent and McHone1988). The original authors’ interpretations of thedepositional setting, age, and architecture of the structurewere insightful for their time but have never beenreassessed in light of post-1960s advances in planetaryscience (Osinski and Pierazzo 2013).

In the original manuscript on the Glasford structure,Buschbach and Ryan (1963) modeled it as a buried pit

filled with breccia, akin to a simple crater. However, animpact structure of this size in a sedimentary target wouldbe expected to exhibit a complex rather than simplemorphology (Osinski and Pierazzo 2013), suggesting thatthe original model is incorrect. Buschbach and Ryan(1963) hinted at a marine origin for at least some of thepostimpact deposits, via reference to graptolite fossils,but did not make this explicit. A reassessment of theGlasford structure, therefore, represents an opportunityto augment the literature on syn- and postimpactsedimentary successions in marine impact craters (Dypvikand Jansa 2003; Azad et al. 2015). Additionally, the ageof the structure has heretofore not been well resolved andhas been variously given as early Cincinnatian (i.e., LateOrdovician; Buschbach and Ryan 1963), ~470 to 485 Ma(i.e., Early Ordovician; Schmieder et al. 2015), or

Meteoritics & Planetary Science 1–24 (2019)

doi: 10.1111/maps.13401

1 © The Meteoritical Society, 2019.

<430 Ma (i.e., Silurian or younger; Hodge 1994; EarthImpact Database 2011). The age question holds particularinterest because the early Cincinnatian estimate implies apossible connection to the “Great Ordovician MeteoriteShower,” a period of enhanced meteorite influx caused bythe breakup of a large asteroid—the L-chondrite parentbody (LCPB)—in the main asteroid belt at approximately470 Ma (Korochantseva et al. 2007). The LCPB breakuphas been described as “the largest documented breakupevent in the asteroid belt during the past 3.5 Gyr” and isthought to be responsible for the disproportionatenumber of Middle to Late Ordovician impact craters onthe Earth (Korochantseva et al. 2007; Orm€o et al. 2014;Schmieder et al. 2015; Bergstr€om et al. 2018; Li and Hsu2018).

In this article, we identify the Glasford structure asa buried marine-target impact crater with a postimpactsedimentary succession broadly similar to those seen atother marine-target impact structures. Additionally, weillustrate and describe features that we interpret asevidence of shock metamorphism, including shattercones and planar features in crater-fill quartz grains; indoing so, we confirm the hypervelocity impact origin ofGlasford, which until now had been ambiguous due toinadequate documentation of previously discovered

shatter cones (McHone et al. 1986; Sargent and McHone1988). Finally, we use biostratigraphic data to date theimpact with much greater precision than in previousstudies. We establish that the age of the Glasfordstructure falls within a few million years of several otherOrdovician impact structures, providing evidence of apossible link to the Ordovician meteorite event.

GEOLOGICAL BACKGROUND

Location and Geological Context

The Glasford structure (hereinafter “Glasford”) islocated in Peoria County, Illinois, USA, at approximately40.60°N, 89.78°W. The center of the structure is 3.5 kmnortheast of the Village of Glasford and 10 km southwestof the suburbs surrounding the larger city of Peoria(Figs. 1A and 1B). The structure is buried beneath 350 mof Ordovician through Pennsylvanian (Upper Carboniferous)rocks and Quaternary materials. It is largely overlain byagricultural fields, with low-relief surface topography thatshows no obvious evidence of the structure (Fig. 2).

Glasford is in the northwestern Illinois Basin(Fig. 1A). Buschbach and Ryan (1963) interpreted thestructure as being immediately overlain by the Upper

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Fig. 1. A) Map of the Illinois Basin showing the locations of the Glasford structure and the cities of Peoria and Chicago. B)Location of the Glasford structure relative to the Village of Glasford in southern Peoria County, Illinois. Dotted line shows anapproximate outline of the structure (after Buschbach and Ryan 1963). Wells that yielded the geological samples or geophysicallogs examined in this study are also shown. Line A–A0 is a trace of the cross section shown in Fig. 7; line B–B0 is a trace of thecross section shown in Fig. 8. Numbers represent section numbers within township T7N, R6E (Public Land Survey Systemcoordinates).

2 C. C. Monson et al.

Ordovician Maquoketa Shale Group (but see furtherdiscussion below). The typical stratigraphic successionbeneath the Maquoketa is predominantly carbonate withlesser sandstone and shale through the Upper Cambrian,underlain by a thick interval of Middle Cambrian Mt.Simon Sandstone that nonconformably overlies igneousbasement rocks (Fig. 3). The Maquoketa Group isunderlain by a succession of Galena Group, PlattevilleGroup, and Joachim Formation carbonates and St.Peter Sandstone. At the Glasford City Well (API121430001800)—the closest well to penetrate to the St.Peter—the Galena–Platteville–Joachim interval is 85 mthick. The thickness of most pre-Platteville units ispoorly constrained in the Glasford area, but regionaltrends suggest a St. Peter thickness of about 60 m andan aggregate thickness of roughly 1010 m for the pre-St.Peter succession, including 460 m for the Mt. Simonalone (Illinois State Geological Survey, unpublisheddata). The estimated total thickness of the sub-Maquoketa-to-Precambrian-basement succession at theGlasford location is therefore approximately 1155 m.

Buschbach and Ryan (1963) proposed that theimpact occurred near the onset of Maquoketa Groupdeposition, during the transition from the limestones ofthe underlying Galena Group to the shaly basal unit ofthe Maquoketa. If correct, this would imply a shallowmarine depositional setting at the time of impact. Themaximum sedimentary load experienced by the structureafter its formation is not well constrained, but probablysignificantly exceeds the current 350 m of overlyingsediment thickness; deposition and subsequent erosion

of thick post-carboniferous sedimentary successions inthe Illinois Basin have been hypothesized based on linesof evidence such as vitrinite reflectance measurementsand clay mineral evolution (Gharrabi and Velde 1995;Rowan et al. 2002).

Previous work

The presence of a structural high at the Glasfordsite was known by the early 1950s, but the feature wasnot recognized as a buried meteorite impact structureuntil the 1960s (Buschbach and Ryan 1963). The firstwell drilled into the structure was the Peters #1 well(API 121430070000; Fig. 1B) in 1952, whichencountered and cored approximately 40 m of atypicalstrata (here interpreted as crater fill) at its base. Duringsubsequent efforts to site a new gas storage field in thearea, a gravity survey and well-drilling program wereconducted, which established the presence of a domewith an estimated diameter of 4 km. The Cowser #1well (API 121430126400; Fig. 1B) was drilled andlogged to a depth of 819 m in the center of the domeand cored to a depth of 762 m. Shallower wells drilledcloser to the rim of the structure yielded additionalgeophysical log coverage, as well as a small amount ofcuttings from strata beneath the typical Maquoketasuccession. During the mid-1960s, the Central IllinoisLight Company commenced natural gas storage invuggy Silurian dolomite overlying the dome (Oborn andRyan 1967), a practice that continues today underAmeren.

Fig. 2. Site of the Glasford structure near the Village of Glasford in Peoria County, Illinois. No surface expression of the structure,which is buried under 350+ m of overburden, is evident. The image was taken at the approximate location of the southern rim,looking north toward the center. Image courtesy of Charles O’Dale.

The Late Ordovician Glasford structure 3

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Fig. 3. Stratigraphic column of the Cambrian–Ordovician section in the Illinois Basin. The Glasford impact likely occurred nearthe end of Platteville Group deposition or the start of Galena Group deposition (see further discussion in Biostratigraphysection), as indicated by the star. Strata affected (uplifted, excavated, and/or damaged) by the impact are shown by a shadedbar. The total depth of impact-related deformation is unknown but may reach Precambrian basement rock. Younger strata andQuaternary materials overlying the Maquoketa are not shown.

4 C. C. Monson et al.

Buschbach and Ryan (1963) published the onlymanuscript-length study of Glasford in the geologicalliterature to date. They described the upper half of thesub-Maquoketa “disturbed section” as dolomite blocksup to several meters thick in a fine breccia matrix andthe lower half as a mixture of intermingled lithologies,including some contorted shale beds (Buschbach andRyan 1963). Although they were able to recognizeblocks of certain Cambrian formations in the centraluplift, they judged the Ordovician Galena, Platteville,and St. Peter units (Fig. 3) to be entirely absent.Identifiable units nearer the base of the cored intervalincluded Eau Claire Formation samples that theyestimated as being elevated some 240–300 m above theirusual stratigraphic level, as well as 60 m of Mt. SimonSandstone (most of it uncored) at the base of the well.They interpreted the structure as a buried pit full ofbrecciated rocks, with a subtle central dome (pierced bythe Cowser core) attributed to postimpact uplift (fig. 5in Buschbach and Ryan 1963). McHone et al. (1986)and Sargent and McHone (1988) subsequently identifiedshatter cones within the Cowser core. Oborn and Ryan(1967) published a manuscript in an engineering journalthat focused on the gas storage field overlying thestructure, while Sargent and McHone (1988) providedfurther discussion of postimpact uplift and the positive(+0.7 milligal) gravity anomaly at Glasford. The onlyresearch on Glasford published during the 21st centuryis Cheng et al. (2018), which (in association with thecurrent investigation) examined meso- and microstructureswithin the central uplift core.

MATERIALS AND METHODS

Samples

The present study utilized geophysical logs, cores,and samples collected during the 1950s and 1960s; nonew material was obtained. Approximately 46 m of 4 in(10 cm) core was available in crater-fill material(beneath the Maquoketa Group) overlying the centraluplift at the Cowser #1 well. In addition, 42 m of 2 in(5 cm) core from the crater-fill succession was availablefrom the Peters #1 well, which is located approximately0.3 km northeast of Cowser #1 and is thought to be onthe flank of the central uplift. Well cuttings wereavailable to a depth of 472 m (encompassingapproximately 100 m of postimpact material) at theHuey #1 well (API 121430126000), which is 1.3 kmnortheast of Cowser #1 and penetrates the annulartrough.

Available material was examined in hand samplesand petrographically to assess rock lithology andmineralogy. Additionally, we drew on unpublished core

and sample studies from Illinois State Geological Survey(ISGS) files. We gratefully acknowledge the work ofISGS employees Charles Collinson and Howard Schwalb(Peters #1 well; Collinson and Schwalb 1953) and T.C.Buschbach (Cowser #1 well; Buschbach 1961), as well asconsulting geologists L. Stone, G.H. Otto, and J.P. Riva(Huey #1 well; Stone et al. 1962), whose descriptionsprovided a foundation for our investigations. The presentwork was undertaken as part of regional geologicalcharacterization efforts for the Illinois Basin—DecaturProject (Finley 2014), with subsequent complementarywork under the Center for Geologic Storage of CO2.

Geophysical Log Analysis

Geophysical logs of early 1950s and 1960s vintagewere available for analysis from six wells. Gamma ray,neutron, dipmeter, resistivity, drilling time, and sonic(spontaneous potential and interval transit time) logtypes were included. The logs were primarily used forcorrelation.

XRD Mineralogy

X-ray powder diffraction (XRD) analysis wasperformed on the whole-rock and clay-particle fractionsof samples from a breccia unit overlying the centraluplift at Cowser #1. Sample preparation initiallyincluded powdering in an agate mortar before obtainingwhole-rock measurements. To prevent possible claymixing, samples prepared for clay-fraction analysis weregently disaggregated rather than powdered. The clayfraction was separated by centrifugation and furtherdisaggregated in an ultrasonic bath. To ensure uniformcation exchange, clay fractions separated by centrifugationwere saturated with 10 mL of an approximately 4 MMgCl2 solution. Suspensions were washed and centrifugedwith distilled water at least three times to minimize thecontent of free ions. Oriented mounts were prepared byusing a Millipore membrane filter and a vacuum filtrationdevice. Measurements were carried out in air-driedconditions, after ethylene-glycol saturation, and afterheating for 1 h at 300°C. A Rigaku� Miniflex II DesktopXR diffractometer was used for XRD analyses. The sizeof the divergence slit was 1°, and the size of the receivingslit was 1 mm. Measurement parameters comprised a stepscan in the Bragg–Brentano geometry using CuKaradiation (30 kV and 15 mA) with a curved-graphitemonochromator. Sample mounts were scanned at acounting time of 10 s per 0.02° from 3° to 70° and 3° to30° 2h for whole-rock and clay fractions, respectively.

Rigaku PDXL Integrated XRD software (v. 2.7.2.0)was utilized for mineral phase identification, which wasbased on a comparison of experimental spectra with

The Late Ordovician Glasford structure 5

published files from the International Centre forDiffraction Data (1996). Diffraction data weresubsequently analyzed by using the Rietveld refinementapproach (Young 1993) to assess the amounts of themain mineral phases identified in analyzed samples.

EVIDENCE OF HYPERVELOCITY IMPACT

Products of shock metamorphism, such as planardeformation features (PDFs) or shatter cones, areconsidered unique to hypervelocity impacts andtherefore diagnostic of an impact origin for candidatemeteorite craters (French and Koeberl 2010). Only onetype of shock-metamorphic product, shatter cones, hasthus far been reported from Glasford. The shatter conesin question were only briefly discussed by previousworkers, in conference abstracts without figures(McHone et al. 1986; Sargent and McHone 1988; Dietzand McHone 1991). Although Glasford is listed as aconfirmed impact structure in some compendia (Rajmon2009; Earth Impact Database 2011), some would arguethat this designation is premature if the key evidence forhypervelocity impact has not been thoroughlydocumented (Reimold et al. 2014). We therefore soughtto locate, describe, and illustrate PDFs and additionalshatter cones to firmly establish the case for ahypervelocity impact origin of the Glasford structure.

Shatter Cones

McHone et al. (1986) and Sargent and McHone(1988) reported shatter cones from Glasford. Thespecimens occurred as “rare but well-developed shattercones . . . in fractured blocks of unidentified massivebrittle dolomite and also as dolomite clasts within brecciaveins bounded by fractured and contorted sandstones”(McHone et al. 1986). The shatter cone–bearing brecciaveins occurred at 365 m (~1198 ft) depth in the Cowsercore according to Sargent and McHone (1988) and Dietzand McHone (1991). No depth was specified for thecones found in fractured blocks. Unfortunately, theshatter cone specimens from the original work werestolen from the storage facility where they were beinghoused and are no longer available for study (McHone,personal communication).

Additional shatter cone specimens (Figs. 4 and 5)were identified in the Cowser #1 core in the course ofcurrent investigations. The specimens are small (≤2 cm)and not particularly distinct, but they are recognizableas shatter cones based on features that include curvedsurfaces and diverging striations with alternatingpositive and negative relief (French and Koeberl 2010).They occur in silty gray dolomite at 549–550 m depth.The first group of specimens (Fig. 4) sits at the edge of

a length of 4 in core. The shatter-coned surface isbroken and locally effaced, but multiple examples ofdiverging striations radiating from a central apex arevisible, nested within a larger arrangement of variablyoriented striations consistent with the “horsetailing”pattern seen in shatter cones. The cones seen at theright-hand side of the specimen in Fig. 4 arecomparatively well developed (though seemingly brokenalong the top) and have curved surfaces (note the arcuatebases) representing approximately one-third to one half(120–180°) of a complete cone. The most distinct of thecones is about 1 cm tall. The set also includes what maybe several additional, weakly developed specimens, asseen at the left-hand side of Figs. 4A and 4B. These aremore ambiguous, occurring as possible cone apices alongthe edge of the core, with poor development of striations.

A separate, relatively large candidate shatter conesurface (Fig. 5) was found about 0.6 m above the first setof specimens. This specimen is ~2 cm tall and constituteshalf (~180°) of a full cone. The specimen is not as distinctas those shown in Fig. 4 but exhibits a curved surfaceand faint but discernable diverging striations. It issandwiched between two darker layers of a type ofmaterial that was identified as argillaceous siltstone or“ductile shales” by past workers (Buschbach 1961;Buschbach and Ryan 1963; Sargent and McHone 1988)but may in fact be microcataclasite (Cheng et al. 2018).

Planar Microdeformation Features

Planar deformation features have not heretoforebeen documented from Glasford. Grains exhibitingPDFs are commonly found in crater fill (French andKoeberl 2010), so we performed petrographicexamination of samples from a sandy dolomite intervalinterpreted as crater fill in order to look for candidatePDFs. Quartz grains are moderately common in thisinterval, and a substantial percentage of them exhibitedmultiple sets of subplanar to planar microstructures(Fig. 6). Candidate PFs and PDFs were photographedand described.

Natural processes can create several types ofmicrostructures that appear as parallel or subparallelsets of linear features in quartz grains. These includetrue PDFs, planar fractures (PFs), and metamorphicdeformation lamellae or MDLs (French and Koeberl2010). The first of these, PDFs, are universally acceptedas impact diagnostic. The second, PFs, are somewhatcontroversial. They can be created by hypervelocityimpact but can form under other conditions as well,albeit uncommonly. As such, they are rejected as proofof an impact event by some authors (Reimold et al.2014), while others have argued that PFs can beaccepted as evidence of shock metamorphism if they are

6 C. C. Monson et al.

Fig. 4. Shatter cones in dolomite from the Cowser #1 core, 549.6 m (1803 ft) depth. Scale bar = 1 cm. A) Small specimens atthe margin of a 4 in well core segment. Striations and cone-like shapes are visible along the entire upper edge of the core, withbetter developed shatter coning at right. B) Same view as (A), but with superimposed sketch highlighting distinguishable coneson the striated surface. Solid black line = well-defined cone border. Solid white lines = orientations of well-defined striations.Dotted white lines = weakly developed or broken cone border or striation. Dark line running SW-NE on partial shatter cone atright is a fracture, not a striation.

Fig. 5. Weakly developed shatter cone surface from the Cowser #1 core, 549 m (1801 ft) depth. Scale bar = 1 cm. A) Divergingstriations are visible in dolomite layer sandwiched between two shale or microcataclasite layers, on right-hand side of specimen.B) Same view as (A), but with superimposed sketch showing striations and borders of cone shapes.

The Late Ordovician Glasford structure 7

present in multiple sets and accompanied by otherindicators of impact (French et al. 2004, 2018; Frenchand Koeberl 2010). MDLs can resemble PDFs or PFsbut are usually the product of tectonic forces unrelatedto impact (French and Koeberl 2010). In evaluatingpotential PDFs and PFs, careful examination isimportant to distinguish such endogenic features fromshock metamorphic features.

A key method for distinguishing PFs and PDFsfrom features such as MDLs is to measure theirorientations, which are parallel to crystallographic axes

of the host grain, but this requires specialized expertiseand equipment (i.e., a universal stage) that were notavailable to our research group. Fortunately, severalother criteria exist that can be useful in differentiatingPFs and PDFs from other types of microstructures.These include multiplicity (multiple orientations), evenrather than irregular spacing (in the case of PDFs), and“a strictly planar character” (French and Koeberl 2010).Some quartz grains within the crater fill clearlyexhibited these characteristics (Fig. 6). Grains typicallyexhibited two or three obvious, well-developed PF and/

Fig. 6. Candidate planar fractures (PFs) and planar deformation features (PDFs) in sandy dolomite interval. All grains takenfrom Peters #1, 366 m (1200.8 ft) depth. In images (A–C), scale bar = 50 lm; in image (D), scale bar = 25 lm. A) GrainP1200.8-1-1, showing two well-developed PF sets (indicated by longer white lines) and several poorly developed PF sets (shorterwhite lines) as well as a set of possible PDFs (short red lines along SW edge, highlighted by box; see online version for color). B)Grain P1200.8-1-2, showing two relatively well-developed PF sets (longer white lines) and several poorly developed PF sets(shorter white lines), as well as a possible PDF set (short red lines at NE edge, highlighted by box). C) Grain 1200.8-2-1,showing several poorly to moderately developed PF sets (labeled with white lines) and at least one set of candidate PDFs (redlines highlighted by boxes), as well as irregular fractures (e.g., curved lines along N edge of grain). D) Higher magnification viewof the area that is roughly delineated by the white box in (C). Red lines indicate orientation of a candidate PDF set along thegrain margin. The planar fracture within the white circle has possible weakly developed feather features running NE-SW, asindicated by the dotted white line.

8 C. C. Monson et al.

or PDF sets that were complemented by multipleweakly developed sets (Figs. 6A–C). Some examples ofweakly developed sets exhibited relatively minordifferences in orientation from well-developed sets,leaving some doubt as to whether they were genuinelydistinct (see, e.g., the farthest left short white line inFig. 6A as compared with the farthest right long whiteline); in other cases, the distinction was obvious (see theother two shorter lines in Fig. 6A). There was typicallyno offset at intersections between sets, althoughdisplacement on the order of a few lm was sometimespresent (see, e.g., lower right of Fig. 6A). Distinctlyplanar sets were common, although in some cases, theywere interspersed with “sets” that were subparallel toeach other. The latter appeared curved (see uppermostportion of grain in Fig. 6C) and thus cannot beaccepted as legitimate PFs or PDFs (Reimold et al.2014). The relative length of the features varied, withthe most obvious sets spanning nearly the whole grainin some instances while other sets extended across onlya small portion.

Spacing and width are key criteria to differentiatebetween PFs and PDFs. Specifically, PFs are typicallyslightly open and hence wider than PDFs (>3 lm,versus <2 lm for PDFs), as well as more widely spaced(15–20 lm, versus 2–10 lm for PDFs; Ferri�ere andOsinski 2013). By these criteria, most of the planarmicrostructures identified in the Glasford samples are bestconsidered PFs. Dominant sets of parallel microstructureswere as much as 40 lm apart and in some cases werevisibly open, often with inconsistent spacing.

A considerably smaller subset of the planarmicrostructures appeared to be good candidates asactual PDFs. In documenting newly identified PDFsfrom the Decorah structure in northeastern Iowa,French et al. (2018) concluded that the majority ofplanar microstructures in their samples were PFs andnoted that actual PDFs were typically restricted tosmall areas at the margins of grains. Similarly, theplanar microstructures that we consider most defensibleas PDFs were found at or near grain margins (see shortred lines [highlighted by boxes] at SW edge of grain inFig. 6A, NE edge of grain in Fig. 6B, and S portion ofthe grain in Figs. 6C and 6D). Measured examples are≤1 lm thick and are spaced ≤3 lm apart, with averagelength within individual sets ranging from 10 to 50 lm.

“Feather features” are another type of planarmicrostructure that has recently been put forward aspossibly being impact diagnostic (French et al. 2004;Poelchau and Kenkmann 2011). Rare structuressuggestive of poorly developed feather features wereobserved in our samples (see circled area in Fig. 6D),but it was unclear whether they were genuine featherfeatures as opposed to intersections of well-developed

PFs with poorly developed, discontinuous PF and/orPDF sets. Our analysis of microdeformation featuresinvolved only a small number of samples, so our failureto locate unambiguous feather features does notnecessarily mean that they are not present.

Summary of Evidence for Hypervelocity Impact

We have, for the first time, provided fulldocumentation of shatter cones from the Glasfordstructure, thereby validating earlier publications andconfirming the hypervelocity impact origin of thestructure. We also have documented quartz grains fromGlasford which exhibit microstructures that appear tobe consistent with shock metamorphism. A substantialpercentage of quartz grains in the crater-fill intervalcontains multiple sets of microdeformation features thatwe interpret as PFs. Multiplicity of PFs is consideredimpact specific by some workers (French et al. 2004,2018; French and Koeberl 2010), which would makethis an additional line of evidence that—while notnecessarily conclusive—seems best explained by animpact origin. Rarer quartz-grain microstructures thatappear to be PDFs (based on size and spacing) are alsopresent, as well as possible feather features. In theabsence of information on the orientation of thesemicrostructures, this can only be considered a briefinitial description of features that warrant further study.Additional work using universal stage and/or TEM(Goltrant et al. 1991) and SEM (e.g., Hamers andDrury 2011) would be highly desirable to support orreject a PF or PDF identification. Nevertheless, withinthe context of a confirmed impact origin for Glasford, ashock-metamorphic origin appears to be the mostparsimonious explanation for these features.

CRATER ARCHITECTURE

Seismic imagery of the Glasford structure is notavailable, so details of the larger architecture are mostlyinferred. The 4 km diameter estimate is based on agravity survey reported in Buschbach and Ryan (1963).Although the presence of a central uplift cannot becategorically established without seismic data, it isreasonable to infer that Glasford is a complex ratherthan a simple crater for several reasons. First, a 4 kmdiameter is well above the 2 km diameter estimate citedby Osinski and Pierazzo (2013) for the simple-to-complex crater transition in sedimentary targets.Second, Buschbach and Ryan’s (1963) observation of“Cambrian formations uplifted about 1,000 feet abovetheir normal stratigraphic position” in the Cowser #1core is roughly consistent with the amount ofstratigraphic uplift that would be expected in a complex

The Late Ordovician Glasford structure 9

crater of this size, based on the relationship given inMelosh (1989, p. 136) (0.06D1.1 = 276 m = 906 ft, whereD = final crater diameter of 4 km). Finally, lateralvariation in geological samples is more consistent withthe presence of a central uplift surrounded by anannular trough; the Cowser well contains coarse brecciaand megablocks in the upper part of the disturbedsection (and throughout the cored interval), whereascore, cuttings, and geophysical logs from comparabledepths in more peripheral wells indicate the presence offiner grained units interpreted here as crater fill.Consequently, the structure is interpreted as a compleximpact crater with a central uplift surrounded by anannular trough (Fig. 7).

The central uplift is expected to be roughly 0.9 km indiameter, based on Melosh’s (1989) ratio of 1:4.5 for thecentral uplift diameter to the crater-rim diameter. TheCowser #1 core through the central uplift was describedby Buschbach (1961), Buschbach and Ryan (1963), andCheng et al. (2018). Its structure and lithology are thesubject of detailed work in progress and will not bediscussed at length here, but some brief remarks forcontext are appropriate. As noted above, the portion ofthe Cowser #1 core within the central uplift is composeddominantly of carbonate breccias and megablocks, withlesser components of shale and sandstone. Somesamples, particularly in the lower portion of the core, arerecognizable as formations typically found atsignificantly greater depths in this part of the IllinoisBasin; these uplifted samples are found in roughly thecorrect stratigraphic order, albeit with some mixing(Buschbach 1961). Strongly tilted beds, extensivefracturing, and meso- to microscale faulting are evident(Cheng et al. 2018). Carbonate intervals within the coreare cut by thin dark layers that were interpreted asinjected shales by Buschbach and Ryan (1963) but areactually microcataclasites which are locally associatedwith sulfide deposits that are likely the product ofpostimpact hydrothermal activity (Cheng et al. 2018).

The Glasford structure is overlain by anapproximately 60 m thick interval of Maquoketa Groupstrata (a typical thickness for this part of the IllinoisBasin), as well as a 12–21 m thick dolomitic shale

interval referred to as “basal” or “atypical” Maquoketain previous work (e.g., Buschbach and Ryan 1963).Buschbach and Ryan (1963) noted that younger stratathin over the central uplift and argued that this thinningindicates continued uplift for millions of years afterimpact and possibly continuing until the present day,with about 72 m of total relative uplift. Similarpostimpact increases in relief have been observed atother craters (Tsikalas and Faleide 2007).

The Cowser #1 well is interpreted as being at or nearthe peak of the central uplift, whereas the Peters #1 wellis interpreted as being on its flank. As shown in Figs. 1Band 7, other wells more distal to Cowser #1—Huey #1,Howell #1 (API 121430126100), Clinebell #1 (API121430126300), and Galloway #1 (API 121430125800)—are interpreted as piercing the annular trough. Onlygeophysical logs and cuttings are available from theseannular trough wells. Whether any of these wellspenetrate the full thickness of the crater fill is not clearfrom this evidence alone, but the Huey #1 well has somefeatures that suggest it may have reached the crater floor.Loggers’ notes identify a “possible shear zone or fault”or “fault gouge” zone at the top and mid-depth of abasal breccia interval (see further discussion below) atHuey #1, with an associated slickensided shale fragmentin cuttings from the mid-depth sample (~445 m), andrefer to “red and green shale injected into breccia.” Thetop “fault gouge” zone is coincident with a positivetemperature anomaly, which is consistent with openfaulting or fracturing (faulted or fractured intervals takeup wellbore cement, which emits heat as it dries). TheHuey “breccia” interval may therefore represent the topof a structural feature such as a slump terrace (Kenkmannet al. 2014) or the floor of a shallow crater rim, withinjected breccia dikes and clay giving the appearance of afully brecciated interval in cuttings. Alternatively,postimpact deformation (Tsikalas and Faleide 2007) mayhave resulted in faulting within allochthonous breccias.

The ~1155 m pre-Maquoketa thickness greatlyexceeds the predicted depth of excavation duringcomplex impact crater formation, even allowing for thepossibility that—contra Buschbach and Ryan (1963)—the impact significantly preceded Maquoketa deposition

Fig. 7. Interpreted positions of wells relative to the central uplift and crater rim. Well depths are to scale relative to one another,but the depiction of crater morphology is schematic; actual depth to crater floor and base of central uplift is unknown. Note thatUpper Ordovician–Carboniferous and Quaternary beds overlying the buried crater are not shown here. The trace of the crosssection is shown in Fig. 1B.

10 C. C. Monson et al.

(see further discussion below). Excavation to basementrock therefore did not occur. Depth of excavation incomplex impact craters is poorly constrained, butOsinski et al. (2013) suggested an approximateexcavation depth of 0.035Da where Da is apparentcrater diameter. Based on a 4 km diameter, this yieldsan excavation depth of about 140 m, which wouldpenetrate the upper Knox Group if the impacthappened during Galena or Platteville deposition. Thisis compatible with Buschbach and Ryan’s (1963)observation that the Galena and Platteville groups andSt. Peter Sandstone were absent in the Cowser #1 core.It is likely, however, that some disturbance of the rockmay have reached the igneous basement. The depth ofthe “totally damaged zone” (zone of completelyfragmented rock) beneath the original surface at atypical impact crater has been estimated as d = D/4,where D is the crater diameter (Senft and Stewart 2007;Stewart 2011). This would indicate 1000 m of “totaldamage” at Glasford, which would penetrate most ofthe pre-impact sedimentary column, extending well intothe Mt. Simon Sandstone. Presumably, an area ofpartial damage would extend deeper still.

CHARACTERIZATION OF CRATER FILL

Although geological material from Glasford islimited, in combination with the geophysical logs it issufficient to identify characteristic and laterallypersistent sedimentary units within the crater. Materialinterpreted as crater fill includes all the materialsampled beneath the “normal Maquoketa” at Peters #1,as well as the top 46 m of material beneath the “normalMaquoketa” at Cowser #1 and 40–150 m of material atthe annular trough wells.

For convenience, we have named the crater fillinterval the “Kingston Mines” (KM) unit. The KM’supper boundary is the base of the “normalMaquoketa,” and its lower boundary is the crater floor(i.e., the top of the autochthonous breccia layer) or, atCowser #1, the top of the central uplift. Laterally, it isconfined to the impact structure (Fig. 1B). Its eponymis Kingston Mines, Illinois (population 259 in 2000), thenearest settlement that does not already have a geologicunit named after it. This is, for the time being, aninformal designation that does not carry anyimplication as to the specific lithostratigraphic rank(formation or member) of the new unit.

Stratigraphic Units

Figure 8 shows a cross section correlating KM units(as well as the overlying Scales Shale of the MaquoketaGroup) across several wells in the structure. These units

comprise a dolomitic shale interval a dolomite–limestone interval, a sandy dolomite interval, and atleast one breccia interval. The trace of the cross sectionis shown in Fig. 1B. A composite log showinggeophysical signatures, lithology, and interpreteddepositional processes for the crater fill units is given inFig. 9.

Kingston Mines Dolomitic Shale Interval (“AtypicalMaquoketa”)

Beneath the expected ~60 m of Maquoketa Groupstrata at the crater location is a 12–21 m thickdolomitic shale interval referred to by previous workersas “atypical,” “abnormal,” or “basal” Maquoketa. Thisunit is not known to occur outside the impact structureradius. A thin “top Maquoketa” distinct from the“usual upper Maquoketa” was also identified ondrillers’ logs at some wells but will not be discussedhere.

Buschbach (1961) and Buschbach and Ryan (1963)referred to the 333.5–352.9 m interval at Cowser #1 as“atypical” or “basal” Maquoketa. However, thelithology and geophysical log signatures of this intervalare similar to the KM dolomite–limestone interval(DLI) at Peters #1 and the annular trough wells. Assuch, we interpret the “basal Maquoketa” of theseauthors at Cowser #1 as representing the DLI of theKM and have assigned the overlying 12 m to the“atypical Maquoketa.” Moreover, because we do notconsider a Maquoketa age for this interval to bedefinitively established, we have grouped the “atypicalMaquoketa” with the other crater-fill units and shallrefer to it as the dolomitic shale interval (DSI) of theKingston Mines unit.

The gamma-ray signature of the DSI is slightlyhigher and more variable than that of the overlying“usual lower Maquoketa” or Scales Shale (Fig. 9). Thetransition into the DSI from the Scales Shale is alsocommonly marked by a very small drop in resistivity(not visible at the scale shown in Fig. 9, but distinct onthe original geophysical logs). At the Peters #1 andHuey #1 wells, the DSI is dominated by silty shale orsiltstone with dolomitic streaks. The interval alsocontains small amounts of coarse crystalline dolomite(presumably vug filling) and, based on cuttings, veryfine-grained sandstone or dolomitic sandstone at Huey#1.

The DSI is not abundantly fossiliferous but hasyielded some biological remains. At Huey #1, it containsscraps of carbonized material of unknown nature, whichmay include some graptolite fragments. Huey #1 wellcuttings from the top of the DSI (370 m depth) include acorrugated, tubular, pyritized fossil fragment. Themorphology of the object is similar to that of fossilized

The Late Ordovician Glasford structure 11

bromalites (gut contents) observed in the marine fill ofthe Decorah crater (Hawkins et al. 2018), although acoprolite interpretation is also possible.

Kingston Mines Dolomite–Limestone IntervalThe DSI is underlain by the second division of the

KM, the dolomite-limestone interval (DLI). The DLI is21–34 m thick at the sampled wells. The gamma-raysignature of the interval is more variable, but typicallylower than that of the DSI. This variability results frominterbedding of shale with limestone or dolomite beds.A gamma-ray spike, corresponding to a shale bedapproximately 1 m thick at Cowser #1, occurs 3–9 mbelow the top of the DLI (Fig. 9) and is traceableacross the entire structure, including the central uplift.An increase in resistivity occurs within the top fewmeters of the interval. Dipmeter values in the DLI areonly a few degrees.

The DLI is dominantly dolomite or silty dolomitewith lesser amounts of interbedded shale, but transitionsto limestone in the bottom half of the interval. Parts ofthe interval are horizontally laminated, although thelamination is often more apparent at a micro ratherthan a macro scale (Fig. 10A). Other parts of theinterval are massive, although sometimes appearing with

dark mud stringers. At Peters #1, where the thickestcored interval of the DLI (~25 m) was collected, theupper half of the unit comprises interbedded shale, darkbrown to gray argillaceous dolomite, and limestone.The lower part of the unit is dominantly coarsening-upward argillaceous limestone that locally achieves a“sublithographic” (fine-grained and homogeneous)character. Similarly, the DLI atop the central uplift atCowser #1 is dominantly dolomite, often argillaceous,silty, or both and sometimes with shaly or argillaceouspartings. The DLI at Cowser #1 transitions fromdolomite to limestone approximately 8 m above itsbase. The basal 1.8 m is argillaceous siltstone with wavybedding that alternates between coarser and finer grains.Thin fining-upward cycles are also evident lower in theDLI at Peters #1 (Fig. 10A).

The DLI includes a component of coarser materialsuspended within the fine-grained matrix. These floatingclasts include subrounded to rounded quartz grains,feldspar, fossil grains, and submillimeter-scale dolomiteintraclasts (Fig. 10B). Cuttings from Huey #1 indicatethat the DLI contains a subordinate amount of sandydolomite to dolomitic sandstone within some parts of theannular trough. Fossiliferous mudstone to wackestone,with bioclasts suspended in a fine-grained carbonate

Fig. 8. Stratigraphic cross section of wells in the Glasford structure showing interpreted sedimentary units (datum = top of theScales Shale of the Maquoketa Group). The trace of the cross section is shown in Fig. 1B. No gamma-ray log was available forPeters #1. GR = gamma ray; LN = long normal resistivity; NEUT = neutron; SN = short normal resistivity.

12 C. C. Monson et al.

Fig. 9. Composite log showing geophysical signatures, lithology, and interpreted depositional processes for sedimentary units inthe Glasford structure. Geophysical log is based on the Howell #1 well. GR = gamma ray; LN = long normal resistivity;NEUT = neutron; SN = short normal resistivity.

The Late Ordovician Glasford structure 13

matrix, occurs at both Cowser #1 and Peters #1 in bedsvarying in thickness from ~5 to 30 cm. These beds recurthroughout the DLI, becoming less common andtrending more toward mudstone (i.e., lower fossil graincontent) lower in the interval. Fossil content ranges fromsmall, unidentifiable fossil hash to intact shells 3 cmacross. Fossils have commonly (but not always) beendiagenetically replaced by spar. Erosional surfaces occursporadically throughout the interval and are sometimesassociated with shell lags. Limestone interbedded withshale higher in the interval is typically fossil rich withirregular bedding boundaries, sometimes with load casts,indicating event deposition.

Small-scale soft sediment deformation occursthroughout the DLI in the Cowser #1 core. Specificfeatures observed include load casts (Fig. 11A), smallflame structures (Fig. 11B), and a ptygmatic ribbon ofunknown origin (Fig. 11C). In some cases, loadingfeatures are also associated with brittle deformation (seethe thinner set of white laminae beneath the loadingstructure in Fig. 11A). A millimeter-scale “wrinkle” or“elephant skin”-like texture composed of irregular ridgesand grooves (Fig. 11D) occurs in the upper DLI; thislikely represents either a type of loading structure or asurface where shear movement occurred, but it cannot beconfidently interpreted on the basis of the limitedmaterial available. Major fracturing is uncommon in theDLI, but large (5 mm wide, >10 cm long) veins (presumablyfracture filling) with associated pyrite deposits are present.

Although the DLI is not abundantly fossiliferous, itcontains a range of marine fossils apart from thepreviously mentioned shell lags and fossiliferousmudstone to wackestone. We saw no indications that theKM contains exceptional fossil deposits like theabundant eurypterid cuticles found at the slightly olderDecorah crater in Iowa (Briggs et al. 2018). However,fragmentary material tentatively identified as eurypteridremains is present on bedding planes, as well as other

arthropod fragments of indeterminate affinity and scrapsof plants or algae (Lamsdell, personal communication).Trilobites, brachiopods, trepostomate bryozoans, andother marine organisms are also identifiable in core. Theupper DLI at Peters #1 includes packstone layerscontaining small, often pulverized fossils interpreted byprevious workers as representatives of the “depauperatefauna,” which is generally considered characteristic of theMaquoketa Group (Johnson 2014). Finally, the DLIyields fragments of fossilized graptolites, which areextinct colonial organisms (Hemichordata) whoseremains can be used to date the enclosing strata with ahigh degree of precision (Maletz 2017). The suite isdominated by biserial and dendroid-type graptolites.Identifiable graptoloid taxa include Glossograptus cf. G.hincksii (Hopkinson), Orthograptus quadrimucronatus(Hall), cf. Amplexograptus, and Hallograptus mucronatus(Hall).

Although well-defined ichnofossils are uncommon inthe DLI, extensive mottling in some intervals is probablyan indicator of bioturbation. Chondrites occur near thetop of the DLI at Cowser #1, and a possible smallZoophycos occurs toward the middle of the interval.

Large, vug-filling calcite crystals occur near the topof the DLI. Pyrite occurs in several forms, includingvery small crystals and, less commonly, larger (severalmillimeter) crystals, as well as groundmass pieces 1 cmor more across, mottles, and fracture linings. A book-like set of white mica sheets was found in DLI cuttingsat the Huey #1 well.

Kingston Mines Sandy Dolomite IntervalThe sandy dolomite interval (SDI) of the KM can

be divided into upper and lower units. Only cuttingsand well log notes were available for the upper SDI; buton this basis, the layer is composed of sandstone, sandydolomite, and dolomitic sandstone. The lower SDI iscomposed of sandy or silicified dolomite, or both,

Fig. 10. Kingston Mines dolomite–limestone interval. In both images, scale bar = 500 lm. A) Horizontally laminated, fining-upward silty dolomite, Peters #1, 362.4 m, plane-polarized light (PPL). B) Fossils and other coarser material suspended withinfine-grained sediment, Cowser #1, 352.5 m, PPL.

14 C. C. Monson et al.

containing floating clasts and large vugs, some of whichare filled with friable quartz sand. The upper SDIappears to be present only within the annular trough,whereas the lower SDI is also present on the flank(Peters #1) but not the crest (Cowser #1) of the centraluplift.

The transition from the DLI to the SDI isidentifiable by a sharp gamma-ray drop on thegeophysical logs (Figs. 8 and 9), accompanied by a slightdrop in neutron readings (i.e., an increase in porosity).Short- and long-normal resistivity readings increaseslightly and show mild separation. Dipmeter readings inthis interval are much higher and more variable (~20°–50°) than in the DLI and chaotically, near-randomlyoriented. The transition from the upper to the lower SDIis marked by further leftward deflection in neutron logsas well as a minor gamma-ray drop. A conspicuousincrease in resistivity values also occurs, with increasedseparation between the short- and long-normal resistivityreadings.

The upper SDI is ~5 to 26 m thick at the studiedwells. Loggers’ notes at most wells identify this interval

as having “sandstone-like porosity” based on micrologresponses, sometimes with a thin zone of lower porosityat the top of the interval. Based on cuttings from Huey#1, where this interval has lower porosity andpermeability than at other wells, the upper SDI isgenerally fine grained, but with scattered larger grainsand possibly a broad fining-upward trend. Occasionalequant dolomite crystals, particularly toward the baseof the upper SDI, suggest some vugginess (and vug-filling dolomite growth) is present. Local chert andpyrite also occur.

Approximately 3.7 m of the lower SDI was coredimmediately beneath the DLI at Peters #1 (the upperSDI is absent at this well). The cored interval at Peters#1 consists of white to whitish-gray sandy dolomitewith a locally chalky texture. The dolomite containsabundant vugs and vesicles, as well as floating clastsranging from sand grains up to clasts about 5 mm inlength (Fig. 12A). The sand grains are fine to coarse,frosted, and locally concentrated in vugs. Petrographicimagery shows submillimeter- to millimeter-scale grainsfloating in a muddy matrix (Figs. 12B and 12D); grains

Fig. 11. Soft sediment deformation in the dolomite–limestone interval in the Cowser #1 core. For all images, scale bar = 1 cm. A) Loadcast with brittle deformation of underlying laminations, 351.4 m depth. B) Small flame structures, 332 m depth. C) Ptygmatic ribbon ofunknown origin, 350.5 m depth. D) “Wrinkle” structures at 337.7 m depth, with shell along the broken edge of the core sample.

The Late Ordovician Glasford structure 15

include quartz, chert, carbonate clasts, and fossilfragments. Dolomite replacement of dissolved fossilsoccurs in some samples. The SDI does not exhibit anyobvious sedimentary layering or systematic grain sizevariation in the Peters #1 core, although millimeter-thick, discontinuous, stylolitized micritic stringers orlaminations occur. Vug-lining euhedral to subhedraldolomite is frequently visible in hand samples and inthin sections (Fig. 12C). Some vugs are rimmed or filledwith pyrite or other unidentified Fe-sulfide minerals.

The limited extent of geophysical logs within thisinterval at Peters #1, and the lack of core at other wells,introduces some ambiguity in correlating the sandydolomite from the Peters #1 with the SDIs at otherwells. However, the spontaneous potential andresistivity readings for the sandy dolomite at Peters #1appear consistent with lower SDI signatures at theannular trough wells. Only cuttings are available fromthe Huey #1 well, but they indicate a lithology like thatof the lower SDI at Peters #1. Notes on logs from otherannular trough wells suggest that the unit contains somesandstone intervals. Sand in the SDI at Huey #1 islargely very fine grained, although larger, well-roundedloose quartz grains occur in cuttings, and a broadfining-upward trend is apparent (particularly in the

lower part of the unit). The Huey #1 cuttings include asignificant amount of white chert lower in the interval,as well as clusters of equant dolomite rhombssuggesting common vug-filling crystal growth. No fossilswere observed in core from the SDI interval of Peters#1, nor in cuttings from any of the wells that penetratethe unit, but drillers’ notes from Clinebell #1 state thatthe lower SDI contains “rare poorly preserved fossils.”

The SDI ranges in thickness from as little as 15 m(Huey #1, where the lower SDI may be absent) to 66–70 m thick at the two other wells (Howell #1 andClinebell #1), where it was fully penetrated.

Kingston Mines Breccia: Cowser #1As previously mentioned, the SDI is absent at

Cowser #1. Instead, the interval immediately beneaththe DLI and overlying the central uplift (approximately353–364 m depth) consists of polymict breccia that weinterpret as allochthonous. On geophysical logs, thebreccia interval is characterized by a lower gamma-rayreading than the overlying DLI. The base of the interval(i.e., interpreted top of the central uplift) is marked by aleftward deflection on neutron logs and increasedseparation between the short- and long-normalresistivity curves (Fig. 8), as well as a lithologic

Fig. 12. Kingston Mines sandy dolomite interval. In image A, scale bar = 1 cm; in images (B–D), scale bar = 500 lm. A) Vuggycore sample, Peters #1, 363.9 m. B) Grains and fossil fragments in dolomite matrix, Peters #1, 366 m, plane-polarized light(PPL). C) Vug-lining dolomite crystal, Peters #1, 372.1 m, cross-polarized light (XPL). D) Grains in dolomite matrix, Peters #1,366 m, PPL.

16 C. C. Monson et al.

transition to banded dolomite with the occurrence ofslickensides and obvious breccia limited to veins.

The breccia is matrix supported but rich in clasts.Clasts are polymict and range in size from sand grainsize (Fig. 13B) to several centimeters or more(Fig. 13A). Clast distribution is uneven, with someintervals being dominated by finer material, althoughsome of the finer intervals may be large altered blocksof country rock that are not easily recognizable in the4 in (10 cm) core. X-ray diffraction analysis shows thatmineralogy of the upper breccia (353.6–356.9 m depth)at Cowser #1 is dominated by dolomite (Fig. 14), withminor quartz and K-feldspar and localized veins ofpyrite. Sporadic mica laths occur whose weatheringgave rise to mixed layer illite-smectite, which is thedominant component of the clay fraction. In thinsections, the occurrence of dolomite is threefold innature. First, it emerges in the form of a fine-grainedmatrix, especially present at 353.6 m, where rounded tosubrounded quartz grains float in the matrix(Fig. 13B); second, as rhombohedral sand- and silt-sized crystals interspersed throughout the matrix(Fig. 13C); and last, as intraclasts, commonly at themillimeter scale, that also float in the matrix. Localized

fractured zones occur and are commonly stained reddishor contain opaque minerals, likely Fe-sulfides, asindicated by XRD. These fractured intervals form ananastomosing network that results in a microbrecciatedtexture (Fig. 13D). Abundant macroscopic pyrite alsooccurs at the very top of the interval, just below the DLI.Samples lower in the zone (~355 to 357 m) have amarbleized character in both hand samples and thinsections. The breccia largely does not appear to be melt-bearing, although one feature that may be an irregularlyshaped, partially altered impact melt clast oraccretionary lapillus was observed (Fig. 12E; cf. fig. 7.7bin Grieve and Therriault 2013).

Kingston Mines Breccia: Annular Trough WellsNotes on geophysical logs indicate the presence of a

breccia layer beneath the SDI at wells located distally tothe central uplift. This interval was sampled only at theHuey #1 well, where 42 m of cuttings were taken;consequently, its lithology, distribution, and relationshipto the breccia unit at Cowser #1 are poorly understood.The interval has markedly decreased resistivity, as wellas reduced separation between short- and long-normalresistivity curves, relative to the lower SDI (Fig. 8). At

Fig. 13. Cowser #1 breccia interval. In image (A), scale bar = 1 cm; in images (B–D), scale bar = 2 mm; in image (E), scale bar= 500 lm. A) Core sample, 360 m depth. B) Image taken in XPL, 353.6 m, showing intraclasts of dolostone floating in a fine-grained dolomitic-mud matrix. C) Image taken in PPL, 355.7 m. Note the dispersed rhombohedral crystals throughout and theabundant pyrite-filled veins. D) Image taken in PPL, approximately 356.6 m. Note the anastomosing network of pyrite- and claymineral-filled fractures in the predominantly dolomite host rock. E) Possible altered accretionary lapillus or cryptocrystallinemelt rock, 353.8 m, PPL.

The Late Ordovician Glasford structure 17

Howell #1, much of this interval exhibits neutron,interval transit time, and resistivity log readingscomparable to those of the upper SDI, suggesting thatit may have comparable porosity and permeability.

Cuttings from the Huey #1 well, where the brecciaunit begins at approximately 430 m depth, show a rangeof lithological and mineralogical components, includinglimestone and dolomite, siltstone, very fine to fine-grained dolomitic or clay-rich sandstone, abundantchert, and shale or fissile mudstone clasts of variouscolors. More exotic clasts include hard green shalefragments with a waxy texture, potentially fused quartzsand grains, and large well-rounded grains that may beloose, impressed into, or embedded in “shale” clasts.Pyrite is common. Some of the dolomite samples arecoarsely crystalline, suggestive of vug fill, and othercarbonate samples have a bright white color similar tothat of the fracture-fill limestone observed in other partsof the Glasford structure.

DISCUSSION

Depositional Processes

Paleontological and sedimentological characteristics(including marine fauna, soft-sediment deformation, andfloating clasts in muddy matrix) show that the Glasfordstructure is a marine impact crater (or “hydrobleme,”cf. Orm€o and Lindstr€om 2000). Based on sample studiesand comparisons with sedimentary successions at othermarine impact craters (Schieber and Over 2005; Dypvikand Kalleson 2010; Azad et al. 2015), the KingstonMines unit is interpreted as sediments deposited withinthe crater either late synimpact or postimpact.Figure 15 shows an interpretive schematic of thedistribution of KM units within the Glasford structure.

Depositional processes for the KM took place withina well-defined crater. Multiple lines of evidence suggestthat a topographic low existed on the seafloor at the

Fig. 14. Whole-rock X-ray diffraction patterns for samples from the upper breccia at Cowser #1 with representative mineralreflexes identified. Dol = dolomite; Qtz = quartz; Sa = sanidine.

18 C. C. Monson et al.

impact site throughout the entirety of KM depositionand possibly throughout Maquoketa deposition. Theseinclude the presence of an atypical, thin “top”Maquoketa unit (suggesting additional accommodationspace within the crater relative to the surroundingseafloor, even at the end of Maquoketa deposition), aswell as the preservation of soft organic features such asbromalites or coprolites in the upper KM, whichsuggests a restricted low-oxygen setting (Hawkins et al.2018). The KM deposits over the central uplift at Cowser#1 are evidence that the central uplift did not protrudeabove the rim of the crater. However, the presence ofshell lags, bryozoan fragments, and other such fossils inthe DLI at Cowser #1 and Peters #1 suggests that thetop of the central uplift was shallow enough for a photiczone to be present, with oxygenation conditionshospitable to normal marine fauna. Compaction overtime of crater-fill sediments and/or central uplift brecciasmay have contributed to the maintenance of atopographic low.

Dypvik and Kalleson (2010) postulate a generalizedsuccession of syn- and postimpact depositional faciesthat might be expected in a typical marine-target crater.These include (in order of occurrence, and of decreasingdepth in the crater fill) avalanches and slides; slumpsand liquefaction; mass flows (debris, mud, and grain);density, turbidity, and tsunami-derived currents; andfinally, postimpact suspension deposits. Not all of thesefacies would necessarily be expected at every crater, norare they all evident in the available material fromGlasford, but aspects of this succession can be inferred.

The breccia unit at Cowser #1 is interpreted as apossible resurge breccia consisting of allochthonous

material transported by water that rushed into thecrater to fill the void formed by the excavation process(Orm€o and Lindstr€om 2000). Given the extremelylimited information available, it is not possible toconfidently interpret the nature or origin of the brecciaunits encountered in the annular trough wells, but thebasic mechanics of marine crater formation suggest thatthe upper layer of breccia logged at these wells (Figs. 8and 9) may be resurge breccia as well. The deeperbreccia layer logged at some of the annular trough wellscould represent either a separate resurge stage (cf. Orm€oet al. 2007); a synimpact deposit such as an avalancheor slump (Azad et al. 2015); or something akin to the“water-blow” breccia formed at the Lockne crater(Lindstr€om et al. 2008).

The SDI likely represents the mass-flow anddensity-current units in Dypvik and Kalleson’s (2010)generalized succession. The bimodal, unstructurednature of the lower SDI is consistent with debris flowsof unconsolidated limey mud entraining sand grains andlarger clasts. Furthermore, the presence of apparent PFsand PDFs is consistent with a crater-fill identification(French and Koeberl 2010). The upper SDI is difficultto interpret due to the absence of core, but it apparentlytrends coarser than the lower SDI, with sandstone anddolomitic sandstone present as well as sandy dolomite.It may therefore represent the density current/tsunami-derived current stage which follows the mass flow stagein Dypvik and Kalleson (2010).

The determination of the mechanism that createdvuggy porosity in the lower SDI is beyond the scope ofthis study. Vuggy porosity can form due tohydrothermal activity, which is often known to occur

Overlying Strata Dolomitic ShaleInterval

Sandy Dolomite Interval

Dolomite-Limestone Interval

AllochthonousImpact Breccia

soft-sedimentary deformation Autochthonous Impact Breccia & Megablocks

fractured bedrock fract

ured

bed

rock

Fig. 15. Distribution of crater-fill units within the Glasford structure. Schematic not to scale. Changes to structure frompostimpact uplift not shown.

The Late Ordovician Glasford structure 19

after impact events (Kirsim€ae and Osinski 2013), andthe presence of sulfide minerals in these deposits isconsistent with hydrothermal activity. However, vuggyto cavernous porosity and sulfide mineral occurrenceare common in Galena-aged formations within theIllinois Basin because of basin-wide fluid migrationevents, such as the one responsible for MississippiValley Type ore deposits in Illinois (Freiburg et al.2012). As such, it may not be necessary to invokeimpact-related hydrothermal activity to explain thevugginess. Nevertheless, the apparent restriction ofpervasive vugginess to the lower SDI (i.e., more closelyadjacent to the fractured bedrock or breccia) could betaken as an indicator of free flow of fluids along thebreccia or fractured bedrock interface, or both. Themechanism for sandy filling of some of the disconnectedvugs within the lower SDI is also unknown. Onepossibility is that lightly cemented clasts of St. PeterSandstone (see Fig. 3 stratigraphic column) wereentrained and subsequently reduced to loose grains bydiagenetic processes that affected the cement. Thefrosted appearance of sand grains within the SDI(Collinson and Schwalb 1953) is consistent withderivation from St. Peter Sandstone deposits displacedby the impact.

The DLI represents a combination of (primarily)postimpact suspension deposition and (secondarily) densitycurrents. Soft-sediment deformation in the DLI occurspredominantly as loading (i.e., flames and load structures)of water-saturated sediment, which could indicateliquefaction of saturated, unconsolidated sediments orgravity flows of material into the paleotopographic low ofthe crater. Fossiliferous wackestone to packstone appearsto be shell lags. Interbedded shale represents interveningperiods of suspension deposition, as do some of the fine-grained limestone and dolomite deposits. Preservation ofgraptolites is also consistent with suspension deposition ina generally quiet environment punctuated by eventdeposition. Trace fossil evidence supports a restrictedenvironment; Zoophycos is commonly indicative of low-energy subtidal settings, including depressions on theseafloor with restricted circulation and low oxygen levels(MacEachern et al. 2012), consistent with depositionwithin a well-defined crater deeper than the surroundingseafloor. Monospecific Chondrites are also consideredindicative of low-oxygen conditions (but usually deeperwater, i.e., outer shelf).

The DSI appears to be transitional between the restof the KM and the “regular” Maquoketa Group. Itcould represent a return to regular sedimentationconditions near the onset of the transgression thatinitiated Maquoketa deposition, although, as discussedin the next section, it is unclear whether the age of theunderlying KM units is compatible with that scenario.

Biostratigraphy

The graptolite suite in the DLI at Peters #1 isdominated by biserial and dendroid-type graptolites tothe exclusion of Early-Mid Ordovician dichograptids,indicating a Late Ordovician age. The dendroidgraptolites do not provide useful biostratigraphiccontrol, but the others are informative. Although thenumber of specimens is small, the suite is consistentwith the Climacograptus bicornis Zone. The Peters #1DLI graptolite assemblage includes Glossograptus cf. G.hincksii associated with large, marginally preserveddiplograptids of the orthograptid–climacograptid–amplexograptid types but including probableAmplexograptus (based on thecal shape and pattern).This suite is generally restricted to the early LateOrdovician (Sandbian Stage; roughly late Nemagraptusgracilis to early C. bicornis Zone). The DLI at Cowser#1 yielded H. mucronatus (Hall) as well as O.quadrimucronatus. This assemblage would seem to putthe Cowser specimens slightly biostratigraphicallyhigher (younger) than the Peters #1 assemblage; inparticular, O. quadrimucronatus would usually be takento indicate, at the oldest, a late C. bicornis Zone age(Goldman et al. 2007). However, given the close similaritiesin geophysical log signatures and sedimentologybetween the interpreted DLI beds at Cowser and theDLI beds at the other wells, a younger age above the C.bicornis Zone for the Cowser material is unlikely. TheDLI is therefore most likely middle to upper Sandbianin age, or stage slice Sa2 (Bergstr€om et al. 2009),deposited at an indeterminate interval within the C.bicornis Zone. The presence of a seemingly uninterruptedmarine crater-fill succession, tracking the expected stagesof syn- and postimpact deposition, suggests that the ageof the crater fill can serve as a reasonable proxy for theage of the crater itself. The graptolite fossils in the upperKM therefore date the impact event to roughly455 � 2 Ma (Goldman et al. 2007; Ogg et al. 2016).

As previously mentioned, the DLI includes layers offossils (often poorly preserved) that previous workersidentified as “depauperate fauna.” If accepted at facevalue, this identification would usually be taken toimply a Maquoketa age for the deposits in question(Johnson 2014). However, the generally poor conditionof the purported “depauperate fauna” specimens leavesdoubt as to their identity. The Sa2 age indicated by thegraptolite assemblage is roughly equivalent to themiddle Champlainian, suggesting that the impactprobably occurred close to the Platteville–Galenaboundary (Fig. 3). A Platteville age for the KingstonMines would require an explanation as to why theoverlying Galena Group, which is ordinarily 60 m thickin this part of Illinois (Willman and Buschbach 1975), is

20 C. C. Monson et al.

either absent within the crater (if the “abnormalMaquoketa” of Buschbach and Ryan [1963] is indeed amember of the Maquoketa Group) or reduced to 12–15 m thick (if the “abnormal Maquoketa,” whichpostdates the graptolite-dated DLI, is Galena aged). Assuch, we consider it more likely that the impactoccurred relatively early in Galena deposition, but anyattempt to assign the Kingston Mines to a specificformation or group is problematic at present. Additionalbiostratigraphic or chemostratigraphic data would berequired to make a more definitive assessment.

The 455 � 2 Ma age groups the Glasford impactwith several others that have been dated to ~440–470 Ma, including two in the Midwestern United States,Rock Elm (French et al. 2004), and Decorah(Bergstr€om et al. 2018; French et al. 2018). Owing touncertainty in some of these dates, the exact number ofimpact craters that can be attributed to this time periodis open to debate (Orm€o et al. 2014; Schmieder et al.2015; Bergstr€om et al. 2018, 2019; Schmitz et al. 2018;Lindskog and Young 2019). Bergstr€om et al. (2018), ina paper dating the Decorah structure to ~464–467 Ma,listed 15 craters (including Decorah, but not Glasford)that purportedly fall within this time interval. However,Lindskog and Young (2019) argued that several of thesecraters are either “very roughly and/or ambiguouslydated” (Clearwater East, Couture, Hummeln, andDecorah) or demonstrably too young (Slate Islands),implying a more conservative estimate of ~10 craters (notincluding Glasford) which can be confidently attributedto this age range. Thus, the addition of Glasford raisesthe number of known craters formed ~440 to 470 Ma to~11 to 15. Uncertainties aside, even the lower countrepresents an order of magnitude increase overbackground rates. This is in keeping with models(Zappal�a et al. 1998) that predict an enhanced influx oflarger meteors to the Earth within 30 Ma of the ~470 Mabreakup of the LCPB, and craters within this age intervalare often interpreted as being related to that event (e.g.,Schmitz et al. 2018). This assumption is admittedlyuntested for most of the craters in question, andgeochemical evidence suggests that a connection to theLCPB breakup can already be ruled out for one of them(Brent; Goderis et al. 2010; Lindskog and Young 2019).Similarly, in the absence of geochemical evidence of anL-chondrite impactor, it is not possible to establish anunambiguous connection between the Glasford impactevent and the LCPB breakup. However, the age of thestructure is consistent with such a connection.

CONCLUSIONS

Our study is a first step toward reevaluating a long-neglected buried impact structure using a 21st-century

understanding of meteor impact processes and products(Osinski and Pierazzo 2013). Relatively few marineimpact craters are known (in comparison with thecurrent predominance of oceanic cover on the Earth),and the Glasford structure has typically not beenincluded in catalogs thereof (Dypvik and Jansa 2003;Dypvik and Kalleson 2010). Our work adds Glasford tothis data set by establishing a marine origin forpostimpact crater fill deposits (a timely addition, withthe recent identification of an apparent marine crater onMars; Costard et al. 2019). We have informallydesignated these deposits the “Kingston Mines” (KM)unit. The KM bears broad similarities to other marinecrater-fill successions, with subunits that were depositedby resurge, mass flow, density current, and suspensiondepositional processes. Graptolite fossils from the KMwere used to provide a well-constrained, relativelyprecise age for the Glasford structure for the first time.Our 455 � 2 Ma age estimate indicates a possibleconnection between the Glasford impact and the GreatOrdovician Meteorite Shower, raising the tally ofknown impact events to ~11 to 15 within 30 millionyears after the LCPB breakup. Recent work tying theLCPB breakup to the mid-Ordovician ice age andrelated biodiversity shifts (Schmitz et al. 2019)underscores the significance of the breakup and theresulting meteorite shower, and Glasford now adds anadditional data point which may help illuminate thisimportant series of events in Earth’s history.

Although the available core samples andgeophysical data are sufficient to confirm Glasford as aburied impact crater and reconstruct its postimpactsedimentary succession, they provide only tiny windowsinto the structure. Seismic data are the most pressingneed to better resolve the overall architecture of theGlasford structure, but myriad questions regarding theevolution of the structure over time, the history ofhydrothermal circulation, and the nature of theimpactor (among other things) remain to be resolved.Additional work on putative PFs and PDFs, as well asgeochemical tests and drilling of new wells in keylocations, would all be desirable. It is hoped that thecurrent manuscript will provide an impetus for suchwork, and that Glasford will become the focus ofongoing research rather than slipping back intoobscurity for another 50 years.

Acknowledgments—Support was provided to C.C.M.(lithological and stratigraphic study) by the U.S.Department of Energy (DOE) through the NationalEnergy Technology Laboratory (NETL) via theRegional Carbon Sequestration Partnership Programunder Award DE-FC26-05NT42588, under the auspicesof the Midwest Geological Sequestration Consortium

The Late Ordovician Glasford structure 21

(MGSC). Support was provided to C.C.M. (structuralgeology and planar microstructure characterization),D.S., B.S., and G.Z. as part of the Center for GeologicStorage of CO2 (GSCO2), an Energy Frontier ResearchCenter funded by the U.S. DOE, Office of Science,Basic Energy Sciences (BES), under Award DE-SC0C12504. Support was provided for K.B. under aMidwest Undergraduate Research Award, andsupplemental funding was provided by the Departmentof Geology at the University of Illinois at Urbana-Champaign. Staff at the Illinois State GeologicalSurvey, including Daniel Byers, Robert Mumm, SusanKrusemark, Lihang Peng, and Mingyue Yu, providedfigure drafting, logistical assistance, and manuscriptediting. D. King, E. Sturkell, and G. Osinski providedhelpful reviews. We thank James Best, Susan Kieffer,James Lamsdell, and Steve Marshak for providingassistance and ideas. We also wish to thank CharlesO’Dale for granting permission to use his photos of theGlasford site. As part of the university grant program,we acknowledge IHS for use of their Petra software tocreate the cross sections.

Editorial Handling—Dr. Gordon Osinski

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