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A Possible Bentonite Origin for Carbonate
Cone-in-Cone Structures within the Late
Famennian Bedford Shale, Rowan County,
Charles Ewing and Chad Ailes
This paper provides an alternative interpretation for carbonate cone-in-cone structures within Bedford
shale of Rowan County, Kentucky. A literature review of the Bedford Shale depositional environment is
followed by a discussion of the diagenetic process by which bentonite clays are transformed to calcite.
Further evidence is brought forward to support the idea that 0-40 cm thick cone-in-cone-bearing
limestone layers within shales are possibly diagenetic products of K- and/or Ca-bentonite ash beds
originating from yet unknown ultraplinian eruptions. Paleo-environment interpretations are discussed
which suggest the eruptions were possibly hosted out of the forming Acadian range during the
Famennian stage of the Upper Devonian period. Samples and similar cone-in-cone beds adjacent to
bentonite are examined and discussed. Geochemistry is examined which supports seafloor subsurface
reduction of sulfate via bacterial mat biology which then displacively forms crystalline carbonates from
solutes derived from volcanic ash.
Cone-in-cone (CIC) structures are often mistaken as fossils, however, their structure is mostly inorganic
in nature. As their name implies, these structures are made of rock with nested cone or cup shapes. In
cross-section, they appear as concentric cones, stacked together like cups at a coffee shop. In plan-view,
CIC structures may be mistaken for circular burrows or coral without closer inspection.
Cupped structures produce one surface with striated cone features while the other side would display
scaled cone cups (Lugli et al., 2005). Those features are retained even if the sample is completely
silicified (Lugli et al., 2005). After release of pressure, the inner cones can pop out of the lower cups and
become separated by erosion.
Carbonate rocks with CIC structure form lenses and beds from 1 cm to half a meter in thickness, with
individual cones being a millimeter to 20 cm. The apical angle of cones varies from 15-120°, with the
apex often oriented toward the bottom of a layer. In CIC beds, the growth direction is normal to the
bedding plane, with variation of 5-15° (Kolokol'tsev, 2002).
CIC structures are often calcareous in nature, but some have
been found which are siliceous material. Both types include
clays, and may include trace metals, odd carbonate species,
and sulfides, while some even include REEs. CIC patterns are
also evident in gypsum, siderite (Lugli et al., 2005), and coal.
Larger CIC structures are found in the subsurface at bolide
impact sites but have slightly different properties in their cone
CIC structures are hosted in shales (both black shale and
carbonate mud), sandstones, and pedogenic rocks, and are also
commonly associated with septaria and carbonate concretions
(Lugli et al., 2005).
CIC structures are also found in lacustrine limestone (Aso et al.,
1992), especially where there is abundant volcaniclastic
material (as silt or clay) and organic matter. In these cases,
many of the CIC structures have been labeled as stromatolitic in
form. Such lakes have seasonally variable pH values (Aso et al., 1992) and in the referenced case are
from regions of arid paleo-climate.
CIC structures are reported all the way back to Ediacaran rock (Meinhold et al., 2019), and contain the
same recurring structure as found in Neogene freshwater lakes (Geptner et al., 2013).
But what are they?
2. Samples and Methods
A thin carbonate layer was mentioned to exist at an outcrop stop in North Eastern Kentucky. We
successfully located the carbonate layer and collected a hand sample which contained the entire
thickness of the bed and afforded us an approximate 2-inch by 3-inch sample to investigate structures
within the bedding plane.
The sample was examined in the field with a hand lens and found to have an outer rind that was
suggestive of an ash bed, or highly weathered crystalline carbonate. Long, blue-gray, fibrous crystals
were evident on the fresh break side of the sample, even without using a hand lens. The rind had a
yellow-orange stain, indicative of iron-rich components. The sample also felt denser than expected for a
One initial in-field interpretation was that the blue fibrous material may be the result of strontium
content, and the formation of celestite(-ine), a Sr-sulfate mineral. Another option was that the blue
color was imparted by a high inclusion rate of reduced iron.
The sample was cut with a steel and diamond blade, with an attempt to create a profile cut through the
fibrous material. The attempt was completely successful, and we were rewarded with a spectacular view
of textbook-quality cone-in-cone carbonate.
Figure 1: CIC structures in limestone. a) Rygel,
M.C.; b) Wilson, M.A. (Department of Geology,
The College of Wooster)
To help determine the source of minerals and elemental
components of cone-in-cone structures, we employed ED-
XRF for simple atomic analysis. We also used digital
petrographic microscopes with a maximum magnification
of 400x to determine individual grain characteristics.
During scientific examination, we also performed a
literature review of the beds we noted at the outcrop. This
review included details about the formation of the beds
and their deposition environments. We included review of
material which referenced adjacent beds, and the extents
and stratigraphic correlation with other states.
In addition to stratigraphic literature review, we also
included literature pertaining to CIC structures, their
relatives, and their precursor materials. That literature also
provided us with recent information about biologic and
geophysical studies surrounding these structures and their
relatives, which led to an extensive literature search
pertaining to clay mineral diagenesis, and ultimately to
volcanic ash beds.
2.1. Site Location
The sample was collected at an outcrop along KY-801, as
marked with a green dot [Fig. 2], north of Farmers, Rowan
County, Kentucky (marked with red pin).
Based on collection site
images, we determine
that the sample is taken
from the contact of the
Sunbury shale and
Bedford Shale. The
member of the Borden
formation can be seen in
the top of Figure 3b. At
the bottom of Figure 3b
can be seen debris
covering the outcrop of
the Bedford Shale. The
sampled CIC layer is
visible as the orange-
stained base of the
Figure 2: Google Maps images showing location of
outcrop (green dot) north of Farmers, Kentucky
Figure 3: a) Outcrop photo showing
scale of CIC layer. Hammer head is 7”
(~17.5 cm) by 1” (~2.5 cm); b) Outcrop
photo showing (from top down)
Farmers Sandstone Member of Borden
Formation, Sunbury Shale, and Bedford
Sunbury Shale, just above the debris slope.
Figure 3a shows a closer picture of the CIC
layer, with hammer for scale.
A simple surface geology map based on data
(Murphy and Petersen, 2005) from the
Kentucky Geologic Survey [Fig. 4] was used to
determine geologic units at the outcrop.
Geologic unit colors were updated to match
those adopted by the Geological Society of
America (Cohen et al, 2012; Gradstein et al,
2012). Based on photos taken at the location,
the data provided by Murphy and Petersen
(2005) appears to be offset many meters to
the west, as the sample was clearly not taken
from within the Farmers sandstone, and the
underlying shale bed is clearly not a minor
unit of the Farmers member.
2.2. Sample Description
The sample is identifiable as CIC by its
distinctive conical appearance [Fig. 6].
Concentric cones nest inside each other. This
feature was barely notable from the initial sample, but after cutting the sample along its longest axis,
parallel to fibrous grains seen at the initial break surface, that cone structure became very apparent.
The composition of the cones at hand-sample scale is visually that of carbonate, based on rhombohedral
and fibrous crystalline shapes. The
spaces between cones are filled with
darker material, which often appears
rusty orange-yellow on the weathered
outer surface of the sample. Individual
grains of dark material are not
identifiable at the hand-sample scale.
Within the crystalline body, individual
grains of pyritized peloids can be seen.
It was not immediately clear if these
were microscopic pyrite concretions or
pyritized peloids, but reflected light
microscopy shows them as having no
crystalline form related to pyrite [Fig. 5].
Instead, all peloids examined appear to
include possible foraminifera debris.
Two locations of the peloid in Figure 5
appear to show foraminifera necks.
Figure 4: Surface geology of outcrop site.
Figure 5: Pyritized peloid. Image taken at 400x magnification (full screen
shown), with cross-polarized light. Additional reflected light is added from
above to highlight peloid surface. Image was modified for clarity (Contrast:
+50%; Brightness: -50%).
Individual foraminifera species cannot be determined due to the low quality of image delivered by
simple petrographic microscopy.
The outer rind of the sample appears to be heavily oxidized and pyritized. When wet, the sample has a
distinct blue-gray color, however, when dry, the innermost part of the cut sample is stark white in the
calcite mineral region, and the blue-gray fades to a pale yellow-white. Note in Figure 6 the possible
weathering front which surrounds the stark white core (denoted by white arrows).
2.3. ED-XRF Data
Using the Bruker Tracer IV SD-Turbo, equipped with a Silicon drift detector (SDD) and Rhodium target
element, ED-XRF data was collected from the outer rind of the sample [Fig.7a], as well as from two
points within the cut surface of the sample. Cut-face sample regions were chosen to reflect 1) the
cleanest carbonate-only data we could expect from the sample [Fig.7b]; and 2) the darkest trapped non-
peloid grain within the sample [Fig.7c].
To maximize device output, and provide the clearest elemental peaks, the device was run with 40 kV
and 15 µA input. No filter was used, and the device was flushed with helium to prevent atmospheric
attenuation. Scans were collected for 60 seconds each. Figure 7 shows collected data in photon counts
per section (cps).
In the rind sample point, there is a notable Yttrium signature, and a strong signature of both Iron and
Zinc with Sulfur. Silicon and Aluminum are also present, as expected from a contact with shale. Other
trace elements include Manganese and Titanium. Calcium registers lower, even though the sample
being scanned is primarily calcium carbonate. This is due to the earlier absorption effects of Iron and
Zinc in the rind.
Figure 6: CIC photo after being cut perpendicular to bedding plane.
0 5 10 15 20
- keV -
x 1E3 cps
Al Si P S Cl K Ca Ti Mn Fe Ni Zn Sr Zr Rb Y Cu
0 5 10 15 20
- keV -
2.5 x 1E3 cps
Al Si P S Cl K Ca Ti Mn Fe Ni Zn Sr Zr Rb Y Cu
0 5 10 15 20
- keV -
2.5 x 1E3 cps
Al Si P S Cl K Ca Ti Mn Fe Ni Zn Sr Zr Rb Y Cu
Figure 7: ED-XRF
showing a) sample rind;
b) cut face purest
carbonate location; and
c) cut face darkest non-
peloid grain. Scalebar is
shown in kcps. X-axis
shows positions in keV.
Ca-intensity: ~8.151 kcps
Ca-intensity: ~5.824 kcps
In the stark white sample point, abundant clean calcium carbonate is shown. Other components are
significantly lower, with Zinc and Yttrium being almost nonexistent.
The sample point taken for trapped darker material shows a mixed signature of both the carbonate and
the rind. Silicon counts for the dark spot are almost exactly those of the outer rind, however Aluminum
is substantially lower. The dark spot mirrors a portion of the rind’s Yttrium and Zinc signatures, as well as
the majority of its Iron signature.
Figure 8 shows a logarithmic scale in which all three scans are compared. There are a few peaks that
must be noted as false. First, the Na-peak represents a Zinc escape peak as it interacts with the SDD.
Additionally, the Ni-peak is partially-to-wholly false due to Ca-pileup where two Ca-induced photons
strike the SDD at the same time. It is interesting to note that in the weathered rind, a substantial part of
the Ni-peak may be true, especially considering the lower Ca-peak. An escape peak can also be seen
near the marked P-peak in which Calcium is interacting with the SDD.
There is a copper peak shown in the histogram which may represent a combination of true Copper and
the interaction of Zinc and Fluorine in the rind. This ED-XRF device does not have the ability to directly
detect Fluorine due to the lower detection limit of the device, however, there are clear Fluorine-
interaction peaks behind Calcium, Manganese, Iron, and Zinc. Even the partially false Nickel signature in
the rind scan has a small Fluorine-interaction peak. Such Fluorine-interaction peaks are marked in Fig. 8
and are always shown 0.68 keV to the left of each element Fluorine is interacting with.
In this histogram, also note the relationship between Yttrium and Zircon in the dark spot and rind. The
rind is substantially elevated in Yttrium relative to included clay particles within the calcite.
0 5 10 15 20
- keV -
Al Si P S Cl K Ca Ti Mn Fe Ni Zn Sr Zr Rb Y Cu Ag Pd Cd Rh Na
Figure 8: ED-XRF spectral histogram (logarithmic scale). Red) Stark-white calcite; Green) Dark non-peloid grain; Blue) Weathered
rind. Blue arrows point to suspected to element-element interaction concerning possible Fluorine. Scalebar is shown in kcps. X-
axis shows positions in keV.
Figure 9: Petrographic images. a) peloids; b) clay materials trapped in annual rings; c) peloids under reflected light. All images
represent full-screen captures at x100 magnification. Purple coloration is induced by λ plate. Greenish color is induced by
Figure 10: a) Trapped
clastic grains within the
calcite matrix (XPL,
magnification x400, with λ
plate); b) pyramidal calcite
c) CIC structure (XPL,
magnification x40); d)
long quartz grain (XPL,
with λ plate)
2.4. Petrographic Images
The goal of collecting petrographic images was to confirm the presence of organic matter, pyrite
nodules, and clastic material. Considering the clastic material, we sought to find quartz and clays
trapped in both the annual rings, and “floating” within the displacively grown calcite matrix.
In addition to inclusions, we hoped to find remnant dolomite or aragonite. To help explain the
coloration, we also hoped to find sulfates, such as celestite or gypsum. We also looked for remnant
sanidine feldspar, as well as preserved mica flakes.
Figures 9-11 show various trapped clastic grains. Use of the λ plate significantly highlights the growth
directions of crystalline calcite in the CIC structure. Most clastic grains were identified as quartz. All
quartz grains were rounded and had no clear boundaries. Darker material surrounding quartz gains may
or may not be clay, but the microscope used could not show individual grains. Other prominent grains
included abundant pyritized peloids, again with dark fuzzy halos.
There are a few occasions in which a bubble-like structure exists in a photo within the calcium matrix.
These features are 3D structures, and so may be bubbles in the epoxy of the slide.
Figure 11: General image of sample showing "floating" clastic grains in calcite matrix. No visible sulfates or other carbonate
species are directly evident. (XPL, magnification x100, with λ plate)
No clear indications exist for carbonate spherulites, nor where there any radioactive mineral halos
detected. There were no visible sphalerite grains to help explain the zinc content.
For the most part, we see the exact same structures identified by Maher, Ogata, and Braathen (2017).
Individual cones often have quartz grains at the apex, or instead grow from a pyritized peloid. Grainy
clays and calcite are evident at annual rings. It is important to note that this interpretation initially
provided by Maher, Ogata, and Brathen (2017) may be incorrect, because unless the sample is
specifically cut to show the exact 3D apex of a cone, what we are instead seeing is a grain at a 2D apex in
the cut cone, which is not the nucleation point from where the cone may have grown. The 2D apex in
the cut merely represents the wall of the cone, which may include any number of quartz or organic
3. CIC Structure
Literature review indicates that no worker as of the time of this paper has been able to find CIC forming
in a modern environment, so this type of structure cannot be studied during formation directly. Instead,
we turn to examining microscopic structures and trace mineral content within CIC-bearing beds.
3.1. Causation of CIC Structures
Some educational portals have claimed that workers generally agree that CIC form from the growth of
calcareous fibers, but even that claim is unsubstantiated. Researchers can’t agree if these are
concretions, or if they’re something else, such as sudden crystal growth after pressure changes.
Part of the problem is that workers have found CIC structures as horizontal beds, but also as radiating
patterns in paleosols with the chevron pointing back toward where a root existed before diagenesis
(Aassoumi et al., 1992), called “rhizocretions”. Still others have found CIC structurers as crack-fill in
septarian nodules and other iron-rich concretions, which may be radial in nature, rather than tabular.
The various disagreements point out that we don’t know much about what causes CIC structures, nor do
we know much about the chemistry of environments they may have formed in. Studies of CIC have been
progressing since the 1800s (Lugli et al., 2005), and it may appear that with modern technology, we’re
finally getting to the bottom of this mystery.
Some microscopic study has now been done in which workers found muscovite, sphalerite, and even
pyrite framboids on what appears to be spore capsules (Lester, 2017). These finds are largely in between
cones, nesting on stair-case-like ledges on what are called annular rings within the CIC structure. Studies
such as that one point out that there’s at least some biologic activity at some point in the formation of
Other studies relate them to freshwater stromatolites (Geptner et al., 2013) or stromatoform organic
gels in lakes with abundant volcanic ash (Aso et al., 1992). Geptner et al. (2013) proposed that, due to
the identification of CIC structures in freshwater lakes, CIC structures should be further examined for
signatures of biologic formation.
McBride et al. (2003) found that large cannonball concretions in western North America contained
multiple stages of growth, some of which included CIC. Isotopic and thermometric data from those CIC
structures showed that the cannonballs took on high magnesium water at a depth no more than tens of
meters, from which biologic agents probably precipitated the CIC. The variation of isotopic oxygen data
showed that the inner environment was both oxidizing and reducing at different times, and other
isotopic data showed variations in marine and meteoric water supply.
3.2. Existing CIC Hypotheses
One debated problem is whether the cones are formed as primary or secondary structures. Trace
dolomite in CIC structures has suggested that the formation is primary (Hooker and Cartwright, 2018).
Due to inclusions carried by the dolomite itself, that hypothesis suggests that clay particles were not
injected into the CIC structure after growth via fractures, and that the CIC structure doesn’t form over
time, but in a single event.
A long-standing idea is that CIC structures form in association with deep burial and diagenesis, providing
both a temperature and pressure element so that stresses can create this somewhat unique pattern
over time. Such a hypothesis says that burial-induced pressure solution allows for insoluble residues to
exist within the CIC structure.
Another hypothesis for the formation of CIC structures is the reduction of pore pressure, which may
create fractures into which pressurized mineral may migrate (Selles-Martinez, 1994). A more recent
attempt at explanation is that gravity waves propagating through uncompacted shallow-burial
sediments created fractures into which crystal fibers initially formed (Abalos and Elorza, 2011).
Aso et al. (2013) put forth the hypothesis that CIC forms in an organic gel in lacustrine sediments that
favors crystallization due to the high pH of the environment. In that case, CIC is a product of early to
very early diagenesis of volcanogenic clays.
A very detailed summary of existing literature, CIC forms, and subtypes was given by Lugli et al. (2005).
Another detailed summary of concretionary forms and their biogenetic relation was given by Seilacher
3.3. Comparisons with Beef-type CIC
In many cases, the fibrous veins are approximately bedding-parallel, and the so-called concretion – a
term that should be used lightly in most cases of this structure type – may be continuous for kilometers,
or discontinuous over several thousand square km with abundant coeval concretions.
In general, these types produce two forms: Beef (BF), and CIC. BF tends to look like a dry steak in that
the “fibers” are actually more fibrous than crystalline, and so the rock may initially appear similar to
serpentine or asbestos. The CIC variety is unmistakable, except for being initially identified as fossilized
wood. Under microscopy, BF-type is also seen as columnar stacks of tiny cones up to 2 mm in width
BF-type concretions are common in organic-rich shales of marine origin throughout the geologic record
from Cambrian to Paleogene times, with gypsum-type BF concretions being common in evaporite and
lacustrine beds of continental origin. Quartz BF is also common in meta-turbidite assemblages (Cobbold
et al., 2013).
Relief of fluid overpressure can precipitate BF-Type bedding-parallel calcite “veins” within horizontal
fractures caused by that overpressure. Organic matter evolution and clay-mineral dehydration (S-I) are
responsible for the fluid overpressure (Wang et al., 2018). In the referenced study, the calcite veins
occur in isolation or in association with concretions and ash layers (Weger et al., 2019), but had little
direct correlation with ash beds. That study also found that surrounding mudstone and concretion
fracture-fill calcite show little difference isotopically when compared to BF (Weger et al., 2019).
BF-type CIC appears to be an important diagenetic feature of immature black shales, but some studies
have shown that the diagenetic temperature often exceeds the expected geotherm during burial, and
well over the 150°C mark for conversion of Illite to Muscovite (Weger et al., 2019). Could this denote a
hydrothermal signature which modified the local geotherm?
Hendry (2002) also showed that BF-type CIC is related to shallow burial of ferroan and magnesian
minerals which break down into ferroan calcite. That study also showed that the isotopic components of
the BF and adjacent shales was directly related.
Cobbold and Rodrigues (2007) put forth the hypothesis that BF and CIC were the products of available
solutes with tensile hydraulic fracturing and dilatant shear failure respectively. Dilatant shear failure
occurs when a non-Newtonian fluid becomes temporarily stronger and applies a greater normal force
when increasing force is applied to it. This inherent feature is driven by colloidal flocculants. Cobbold
and Rodrigues (2007) grouped these two fracture types as “seepage forces”.
Fibrous BF-type structures have been found which combine from top and bottom and transition directly
into CIC structures (Luan et al., 2019), providing evidence that BF and CIC are perhaps formed in the
same environments under somewhat different localized stress regimes.
BF-type structures may form initially from bicarbonate derived from a mixing of carbon from previous
carbonate dissolution, and organic carbon from both fermentation and thermal decarboxylation (Luan
et al., 2019).
3.4. Carbonate Precipitation Chemistry
Mejri et al. (2012) determined that Magnesium and Sulfate were key players in the precipitation style
and speed of carbonates. In their study, at a fixed temperature, additional magnesium and sulfate
Figure 12: Mejri et al. (2012) depiction of fascicular calcite growth with spherulites in Mg-clay matrix.
decreased crystal growth. When temperature was increased, magnesium induced the formation of
aragonite rather than calcite or vaterite. Increased temperature in the presence of both magnesium and
sulfate led to the incorporation of magnesium ions in the carbonate lattice, with both aragonite and
vaterite being precipitated. Nielsen et al. (2016) also determined that while magnesium alone decreased
crystal growth, addition of sulfate, which probably produced MgSO4 and bound up ions, further reduced
Earth’s marine environment is a brine solution composed primarily of Na+, Cl-, Mg2+, Ca2+, K+, and SO42-.
These same components are common in evaporite settings, in either marine environments or high pH
Pre-salt, alkaline lacustrine settings have been observed precipitating early diagenetic fascicular calcite.
One important instance of such an environment is the northern Campos Basin, Brazil. Lima and De Ros
(2019) documented fascicular calcite growing upward from the clay-water interface, with additional
substrate-hosted calcite spherulites [Fig. 12]. In that setting, the primary sediment is a Mg-clay. Both
growth types displaced the clay. Spherulites where identified as being neoformation of calcite after
diagenesis of Mg-clay. Such fascicular calcite resembles very closely that of CIC structures.
Braissant et al. (2003) showed that the exopolysaccarides and amino acids secreted by bacteria are
responsible for the formation of vaterite and calcite spherulites. Images of vaterite spherules appear
very similar to the growth habit of CIC, except on a much smaller scale. This is important because
exopolysaccarides are what stromatolites use to precipitate aragonite domes, and CIC structures closely
resemble stromatoform domes in many cases.
3.5. Summary of CIC Hypothesis Agreement
Minerals responsible for the formation of both BF-type and CIC-type concretions are thought to be
locally sourced (Cobbold et al., 2013), probably from seawater (Meng et al., 2017). In some cases, CIC
structures appear to form in the process of dedolomitization, in which Mg is removed from dolomite to
produce calcite and periclase (Kowal-Linka, 2010). In that case, vein widening is thought to be from the
force of calcite crystallization. That observation coincides with a previous hypothesis in which CIC
structures formed after the conversion of aragonite to calcite, which requires a volumetric change. In
these hypotheses, it is thought that displacive growth can incorporate clay minerals from adjacent
material into the CIC structure (Fairbridge & Rampino, 2003). Additional displacive growth and
cementation may follow the generation of CH4 and CO2 by sediment-hosted bacteria in the
methanogenic zone (Meng et al., 2017).
Overall, the more popular and widely used explanation for CIC formation is some variation on displacive
crystal growth, usually within a partially consolidated sediment, and with some studies suggesting that
the crystal growth starts in a marine environment as aragonite, displacing clays and eventually being
replaced by calcite (Gillman & Metzger, 1967). Substantial framboidal pyrite in the sediments
represents an early phase of microbial sulfate reduction, which may have induced initial calcite
mineralization (Maher et al., 2015). The presence of skeletal fragments in some CIC and BF provide a
control on the special distribution of calcite nucleation points (Meng et al., 2017).
In recent literature, evidence points to:
1. Some amount of burial is expected, although it can be shallow burial
2. A marine environment is not required, as BF and CIC have both been observed in non-marine
3. Crystal growth is not gravity driven, meaning CIC chevrons do not indicate an up direction, and that
BF-type concretions grow fibers in various directions
4. Widespread CIC formations do not immediately coincide with mass extinctions or extinction-level
5. CIC may form in high pressure and temperature settings, but do not require such
6. CIC may be biologically active at periods in its formation (Maher et al., 2015; Kershaw & Guo, 2016;
Tribovillard et al., 2018)
7. CIC can retain significant information about shale diagenesis and organic matter evolution (Luan et
8. CIC can be a product of early diagenesis (Tribovillard et al., 2018)
4. Geologic Setting
In this section, we examine literature containing references to the Bedford Shale, from which the sample
was taken. In addition, we also examine the extents and transitional forms taken on by various parts of
the Bedford Shale. Information is also examined which gives details about the deposition environment
of the Bedford, and it’s depositional neighbor beds.
4.1. Bedford Shale
The Bedford Shale is a Late Famennian geologic formation (Pepper et al., 1954) and is the basal member
of the Wavery Group (Collins, 1979), which also includes the Maxville Limestone, Logan fm., Cuyahoga
fm., Sunbury Shale, and Berea Sandstone.
The extents of the Bedford shale cross state lines, being represented in Ohio, Michigan, Pennsylvania,
Kentucky, West Virginia, and Virginia (de Witt & McGrew, 1979; Pepper et al., 1954; Gutschick &
At the site location, the Bedford Shale underlies the Sunbury Shale, and overlies the Cleveland member
of the Ohio Shale (Rice et al., 1979), while in many other locations, the Bedford Shale underlies or
fingers into the Berea Sandstone above.
4.1.1. Bedford Shale Age
Interpretation puts the Bedford Shale at between 365 and 358.9 Mya by the current definition of the
Devonian Period (Janssens & de Witt, 1976; Conkin et al., 1980; Pashin & Ettensohn, 1995). However,
there was some conflict as to whether the Bedford Shale belonged in the lowest Carboniferous Period
based on its basal unit (Pepper et al., 1954; de Witt, 1970; Collins, 1979; Dick & Shakoor, 1995), or
whether it straddled the Devonian-Mississippian boundary (Ettensohn et al., 1977; Sable, 1979; Ells,
1979; Edmunds et al., 1979; Cohee, 1979).
Conodont investigation suggests that the upper Devonian date is the best interpretation, as no lower
Carboniferous conodonts were found in either the Bedford Shale, or the adjacent Berea Sandstone
(Gutschick & Sandberg, 1991; Catacosinos & Daniels, 1991). The rock unit is now placed in the Late
Famennian stage of the Upper Devonian, rather than the Kinderhookian stage of the Lower
4.1.2. Bedford Shale Description
The Bedford Shale fm. was first identified in 1870 at Tinkers Creek, Cuyahoga County, near Bedford
Ohio. At that location, the Bedford Shale is a blue-grey shale 26 m thick (Wilmarth, 1938). In other
locations, the Bedford fm. is red clay shale that grades to blue-grey at its base (Collins, 1979). The clay is
often silty or sandy, with hard siltstone sub-members (Cohee, 1979; Harrell et al, 1991; Lee, 1991).
Because of color variations, the Bedford is subdivided into grey Bedford and red Bedford with the Red
Bedford thinning and grading into the grey Bedford south of Columbus, Ohio, along its other margins
(Pepper et al., 1954), and is listed as grey to blue-grey in Michigan (Ells, 1979; Cohee, 1979; Harrell et al.,
1991). In Kentucky, the Bedford Shale appears greenish-gray to gray and contains calcareous
concretions and pyrite nodules (McDowell, 1983).
The formation includes siltstone beds up to 7.6 cm thick which display ripple marks, and which are
interbedded near the bottom (Pepper et al., 1954). In other locations, the Bedford fm. includes
sandstone which may or may not be related to the Berea Sandstone (Collins, 1979; Pepper et al., 1954).
Calcareous siltstone layers can be found within the Bedford with approximate thicknesses of 5 to 10 cm.
Within Ohio, siltstone beds increase in number and thickness in its upper portions, especially south of
Columbus. In addition, blebs of calcium carbonate, marcasite, and pyrite occur in the upper third of the
Bedford (Pepper et al., 1954).
4.1.3. Bedford-Cleveland Contact
Fine-grained silty mudstone laminae can be found toward the bottom of the Bedford fm. (Lee, 1991;
Pepper et al., 1954). Within Kentucky, a transition zone ranging in thickness from a few cm to 1.3 m
Figure 13: Earth 360 Ma, as depicted by Colorado Plateau Geosystems (2016). The position of Bedford Shale deposition is
occasionally appears at the base of the
Bedford, although through most of the rest
of the extent, the boundary with the
Cleveland member of the Ohio Shale is
clear (Collins, 1979; Pepper et al., 1954).
4.1.4. Bedford-Berea Contact
The upper boundary with Berea Sandstone
in Michigan exhibits sand and silt stringers,
indicating erosion of the Bedford shale and
mixing with the sand which was to become
the Berea Sandstone (Ells, 1979). Within
Ohio, the upper contact is extremely
irregular (Pepper et al., 1954) [Fig. 15]. In
those places, Berea sandstone filled the
missing portions (Milici & Swezey, 2006).
Berea Sandstone is not evident at the site
location in Farmers, Kentucky. Instead, the
Bedford is overlain by the Sunbury Shale
(Elam, 1981), but the Bedford had been
previously termed the Berea-Bedford
Sequence with eastern Kentucky (Provo,
1976), due to the sand and silt stringers. In
far north-eastern Kentucky, the Bedford
fm. is called “Berea Sand” (de Witt &
McGrew, 1979). This sandy-silty upper
section is evidence of inter-tongued
Berea-Bedford siltstone (Sable, 1979).
Pre-Berea erosion of the Bedford Shale
removed large sections of the Ohio Red
Bedford, and also cut into the underlying
Cleveland Shale in the southwest. In
southern Ohio, the Berea and Bedford
grade into each other without an obvious hard transition (Pepper et al.,
1954). The same is true of Bedford-Berea transition in Michigan (Gutschick &
4.1.5. Bedford Shale in Kentucky
Bedford Shale is present throughout much of eastern Kentucky (Ryder et al., 2015; Rice et al., 1979) and
is thickest in north Lewis County, thinning and pinching out in Bath, Estill, and Pike counties (McDowell,
1983; deWitt & McGrew, 1979) to the south and west [Fig. 14]. Thickness of the Bedford fm. is about 30
meters in Letcher and Pike counties of Kentucky (McDowell, 1983), which is likely due to proximity to
the Famennian deltas (Pepper et al., 1954). The general shape is a wedge with a narrow edge in the
southwest (Sable, 1979).
Figure 14: Stratigraphic column of Kentucky depicting the Devonian-
Carboniferous transition. Included are important members of the Late
Devonian and Mississippian deltas. Image cropped after Ettensohn et al.
Figure 15: Berea Sandstone
fills erosional channels in
Bedford and Cleveland Shales,
after Pepper et al. (1954).
The topmost section of Bedford in most of Kentucky is a soft blue-gray clay
(Pepper et al., 1954). At the location of this site, Berea sandstone is
missing, and has pinched out in the adjacent Bath County (McDowell,
1983). What remains of the Bedford at the location of Farmers, Kentucky, is
largely clay with abundant oxide hydroxides of iron, intermixed with shale
fragments. Ettensohn et al. (2012) identified an upper CIC layer just under
the hard transition to Sunbury Shale [Fig. 16].
The Bedford Shale is largely fossil-free, except for the northeast basal
siltstone unit, which is topped with invertebrate fossils (Pepper et al.,
1954). Where Bedford Shale is in contact with the overlying black Sunbury
Shale, brachiopod and conodont fossils can be found in a pyritized region
4.2. Famennian Deltas
During the Famennian stage of the Upper Devonian period, Bedford shale
was deposited in Kentucky by several small rivers in a deltaic pattern, with
deposited occurring from east and southeast. During deposition, the
environment was mild, with abundant plant life growing along tributaries.
Equatorial monsoons swept the region from west to east (Pepper et al.,
Much of what is now Kentucky was covered by the shallow Rheic Ocean
(Frazier & Schwimmer, 1987). At the end of the Acadian orogeny, in the
middle Devonian, the Acadian Mountains began to erode. As sea levels also
subsided, material from both the Canadian Shield and the Acadian
mountains were deposited as river deltas prograding into the Rheaic
Ocean, creating the Bedford Shale, and Berea Sandstone (Milici & Swezey,
2006). This deltaic system produced fans all along the northern and eastern
rims of the Ohio Bay of the Rheic Ocean, which spanned multiple states,
including most of Kentucky, Michigan, Ohio, western Pennsylvania, Virginia,
and West Virginia (Pepper et al., 1954).
When the Bedford Shale was beginning to be laid down, anoxic seas were
once again oxygenating, probably due to precipitation captured by the
Acadian range. Initial coarse material influx created the basal siltstone units
of the Bedford Shale in which fossil brachiopods and other fauna are found.
Later increased influx of fine materials made the deltaic system
inhospitable, and so little to no fossils are found in the majority of the Bedford Shale (Pepper et al.,
1954; Rice et al., 1979).
As deposition occurred, sea levels once again rose, reaching median depth at about the time half the
Bedford Shale had been deposited. The sea then fluctuated greatly while the rest of the Bedford was
deposited (Pepper et al., 1954). Initial siltstone beds retain oscillation ripple marks which indicate that
the bed was laid down in a subaqueous environment, whereas the red Bedford Shale in the center of
Figure 16: Stratigraphic
column from Ettensohn et al.
(2012) which shows (in very
fine print) a CIC layer directly
below the Sunbury contact.
the Ohio Bay was deposited subaerially (Pepper et al., 1954). Red Bedford subaerial deposition is also
supported by braided paleo-streams.
Deltaic shale of the Bedford fm. exhibits crossbedding and erosional channels filled with later Berea
Sandstone, with both indicate the depositional environment was shallow to subaerial at times (Sable,
1979). These channels are also evidence for the paleocurrent directions, indicating an overall westward
and southward flow. Ripple marks in the basal siltstone member indicate a northeast wind source
Other CIC-related research has been done on similar deltaic and shallow lacustrine paleo-environments,
such as that of the Upper Triassic Yanchang Formation, China (Milliken et al, 2016). In that study, the
shale was identified as organic-rich lacustrine tarl: >75% of particles of extrabasinal derivation, including
grains derived from continental weathering and also volcanogenic debris (Milliken, 2015; Milliken et al.,
In this section, we examine various mineralogy one might expect from ashfall and tephra deposits after
5.1 Bentonite Clay
Bentonite is a clay formed by the weathering of volcanic ash. The two most important mineral
constituents are Montmorillonite and Beidellite – both from the Smectite clay group. Physical properties
of Bentonite vary considerably and are partially related to the proportions of its individual mineral
species (Spivey, 1940).
5.1.1. Bentonite Chemistry
The primary component of bentonite is Montmorillonite: (½Ca,Na)(Al,Mg,Fe)4(Si,Al)8O20(OH)4.nH2O.
Sodium bentonite expands when wet, absorbing as much as several times its dry mass in water. Calcium
bentonite is a useful adsorbent of ions in solution (Lagaly, 1995).
Conversion of smectite to illite, or illitization (S-I), during burial is a common diagenetic process which
can release large amounts of Fe, Mg, Ca, Na, and Si (McHargue & Price, 1982). Smectite is converted to
illite with increasing depth of burial in sedimentary rocks, based on dissolution of smectite and the
neoformation of illite (Nadeau et al., 1985). Potassium fixation together with neoformation appear to be
the primary mechanisms of conversion of smectite to illite (Šucha et al, 1993).
Calculations indicate that the stability of smectite layers may be a function of composition. Smectites
with high ratios of octahedral (Fe + Mg)/Al appear to resist S-I until temperatures high enough to
produce ordering are attained (Hower et al., 1976).
Bentonite clay can reduce ferric iron to ferrous iron and trap sulfides. Ferrous iron in turn destabilizes
montmorillonite (Pedersen et al., 2017; Lanson et al., 2012). A recent study (Kim et al., 2004) proposed
that the rate of the S-I reaction was accelerated by microbial reduction of structural Fe(III) in smectite.
5.1.2. Silica and Calcium Carbonate from Bentonite
Montmorillonite can be forced to trade ionic and included components with carbonates via the
application of Na (Natale & Helmy, 1992). Ca-carbonates will form in place when the exchange of cations
produces a Ca-rich solution which is then not removed from the bed (Guyonnet et al., 2005). This was
demonstrated in synthetic clay liners where NaCl was added to a Ca-bentonite, which then produced Ca-
Released Si and Ca during S-I can be transferred from shales to sandstones to produce quartz
overgrowths and calcite cements at temperatures as low as 60°C (Boles and Franks, 1979). Authigenic
carbonates cement and can replace diatomaceous ooze, volcanic ash and bentonite beds, and clastic
beds (Hein et al., 1979). Carbonate cement is much more commonly observed in sand bodies adjacent to
Ca-rich source rocks. Dissolution of Ca-feldspars by organic acid-rich fluids, together with S-I, provides Ca
and Mg ions, and carbonate cement would then precipitate with a decrease in CO2 concentration (Lai et
Tephra may also be altered to Ca-zeolite minerals instead of clays, producing clay-poor calcite-cemented
layers (Ver Straeten et al., 2012). Kolodeznikov and Stepanov (1986) identified zeolites formed at the
late stage of diagenesis of ash beds and noted that calcium carbonate was one of the adjacent products.
In that study, the zeolites were directly adjacent to salt domes, which could have provided abundant
sodium to catalyze diagenesis.
5.1.3. Ultraplinian Eruptions as a Source of Bentonite
In many cases, individual ash-bentonite beds have a volume greater than that of the Toba eruption.
Some researchers suggest that the ashes were from a volcanic arc that was on the convergent crust
boundary (Haynes, 1994) due to felsic calc-alkalic traces thought to be from a magmatic source, which is
characteristic of volcanism from a continental crust destructive plate margin.
North American ultraplinian (ash volume > 1000 km3) eruptions appear to have been derived from a
volcanic arc or microplate subduction zone along the eastern side of Laurentia (Bergström et al., 2004).
Where the bentonite thickness and grain size are at their maximum are most likely close to the eruption
sites (Haynes, 1994).
A noteworthy tuffaceous bed can be found as the Rockyford Ash (RAZ) in parts of South Dakota and
Nebraska. RAZ forms the basal member of the Sharps Formation and is approximately 2.5 m thick in the
Badlands of South Dakota (McConnell and DiBenedetto, 2012) reaching a maximum thickness of 17
meters (Raymond, 1986). The member is dated as being deposited 30.05 Mya (McConnell and
DiBenedetto, 2012). One possible source of the ash is the Yellowstone Caldera in Western Wyoming.
The Sharps Formation continues with widespread marker beds, freshwater limestones, and clayey
siltstone (McConnell and DiBenedetto, 2012), many either containing ash, or being partially derived
from ash. RAZ has been altered to zeolite with sections frequently containing more than 40% high-
potassium zeolite (Raymond, 1986).
5.2. Illite Clay
Illite is non-expanding clay structurally similar to muscovite, with more Si, Mg, Fe, and H2O, with less Al
and interlayer K. Illite can precipitate from diagenesis of smectite clays.
The formula for illite clay is KyAl4(Si8-y,Aly)O20(OH)4 , usually with 1 < y < 1.5, but always with y < 2.
Because of possible charge imbalance, Ca and Mg can also sometimes substitute for K. Interlayer cations
prevent the entrance of H2O into the structure.
S-I conversion has long been thought to be caused by abiotic factors: temperature, pressure, time, pH
and the activity of K+ (Ahn and Peacor, 1989; Hong et al., 2017).
Illite crystallinity has been used as an indicator of metamorphic grade in clay-bearing rocks
metamorphosed under conditions between diagenesis and low-grade metamorphism (Frey, 1999). With
increasing temperature, illite can be transformed to muscovite (Gharrabi et al., 1998). Specifically, K-
bentonite taken to a temperature of ~150° C converts to muscovite (Šucha et al, 1993).
K-bentonite is a potassium-rich illitic clay formed from alteration of volcanic ash, and which contains
high I/S ratios (Hong et al., 2017). Altered ashes may have originated via felsic volcanism (Hong et al.,
2017) which can be determined via immobile REE plots. A review of known global tephra k-bentonites is
given by (Huff, 2015). Eighty or more k-bentonite beds have been identified in just the Appalachian
foreland basin (Ver Straeten, 2016).
5.2.2. K-bentonite Alteration
Depositional environments have great influences on the alteration of ashes (Hong et al., 2017). 87Sr/86Sr
and 143Nd/144Nd isotopic composition, clay mineralogy, and trace element geochemistry can help
determine depositional environment (Hong et al., 2017).
Ashes in deep water environments have experienced more intensive weathering (Hong et al., 2017).
Marine environments promote early diagenesis via a process called halmyrolysis (Günal-Türkmenoğlu et
al., 2015) in which ferromagnesian minerals convert to glauconite. This process produces small-grain
potassium-iron phyllosilicates, which can then accumulate more potassium, as well as Rb.
During diagenesis, by homogenization of tephra with marine and pore water, volcanic glass
progressively alters or transforms into smectite, then mixed-layer illite-smectite, and finally illite (Günal-
Türkmenoğlu et al., 2015). Element ratios can also be helpful in determining intensive chemical leaching
in deep-water environments. Those plots use Si/Al, K/Al, Y/Sr, and Zr/Sr (Hong et al., 2017).
Open-water reducing conditions promote neoformation of Mg-smectite. This paragenesis can then be
used as indicator of shallow alkaline lake and playa environments (Hay et al., 1991). Meteoric water
infiltrating a thick tuff layer becomes more alkaline and saline with depth (Gottardi, 1989). High-
salinity water was also observed to assist bentonite and illite to settle faster (Liu et al, 2017). Other
factors of ashfall preservation include: greater marine body depth, low biologic disturbance, and high
sedimentation rates (Ver Straeten et al., 2012).
Surface glass is altered to clay minerals while deeper glass is converted to zeolites (clinoptilolite,
chabazite, phillipsite). In the pH range 9-10, up to 70% of the sediment from ashfall is recrystallized
(Hay et al., 1991). According to stable isotope evidence, playa oxidizing conditions (dry lakes, salt flats,
and sabkha) favor precipitation of Fe-illite and K-feldspar. In experiments by Small (1994, 1995),
carbonate and organic anions buffered the activity of potassium, and the rate of illitization increases by
several orders of magnitude.
It is important to note that K-bentonite beds may represent multiple ashfall events (Ver Straeten , 2016)
in what appears to be a single bed. It is therefore important to identify separating planes, such as zonal
clay mineralogy, and to collect and date multiple samples from the full vertical range of a bed. Such beds
may also include fossil layers at transitional lamellae, as well as glauconite and phosphate nodules (Ver
Straeten, 2016). Other evidence may include changes to grain size and orientation of volcanic grains
Volcanic ash that falls into marine settings commonly alters to smectite deposits, whereas volcanic ash
falling into non-marine environments may alter to kaolinite deposits. Kaolinite may then lithify as a
claystone known as a tonstein (Bohor and Triplehorn, 1993). The relationship between tonstein and
volcanic ash was first made by Price and Duff (1969).
Tonsteins are common in coal-forming environments due to the preservation of thin ashfall deposits by
abundant organic matter. The thin ash beds contain kaolinite, quartz, zircon, and biotite, as well as
partially kaolinitized biotite (Weiss et al., 1992), and may also include the high temperature K-feldspar,
sanidine, Ca-amphibole, and Bytownite, the Ca-endmember of plagioclase.
In addition, forest and peat bog settings have low detrital input due to the baffling effect of organic
matter (Bohor and Triplehorn, 1993). Low inclusion of detrital illite and quartz indicate rapid deposition
into aqueous environments (Lyons et al., 1992). The highly leaching organic acid environment is partly
responsible for the alteration of ashfall beds to kaolinite (Bohor and Triplehorn, 1993).
Bed thickness also plays a part in alteration. Thick beds may show zonation of clay mineralogy (Bohor
and Triplehorn, 1993).
Tonsteins have been paired with bentonite beds for over a century to determine stratigraphic
correlation of coal seams (Spears, 2010). In one case, a tonstein and bentonite bed showed a clear
weathering front where smectite in the bentonite was weathering directly to kaolinite in the tonstein
(Senkayi et al., 1984).
Zhou et al. (1994) determined that zircon characteristics in tonsteins could be related to the distance
from the volcanic source, and that grain size changes along the tonstein also identified the direction of
5.2.4. Use of Ash Beds as Geochronologic Indicators
Ash beds are not directly associated with every notable faunal extinction event or any significant
lithologic change in the sedimentary record (Bergström et al., 2004), however, there is some correlation
between ash beds and extinction events of various scales.
Ultraplinian Ordovician Millbrig and Kinnekulle ashbeds were correlated by date and thickness
(Bergström et al., 2004; Mitchell et al., 2004) to reconstruct a possible location for an island arc
explosive eruption source [Fig. 17]. Data was combined with geochemical and conodont
biostratigraphical evidence to support the correlation.
Using the 40Ar/39Ar method, diagenetic clays can be dated (Dong et al., 1997). Using transmission
electron microscope (TEM) with vacuum-encapsulation laser techniques allows for the separation of the
timing of metamorphic growth and cooling from the time of diagenesis (Dong et al., 1997).
Israelson et al. (1996) showed that some carbonate concretions had elevated U/Pb ratios, which allowed
for precise geochronology. Those same concretions yielded younger dates based on material taken from
their internal cone-in-cone structures, which Israelson et al. concluded provides the date of final stages
of sediment compaction.
5.2.5. Use of CIC Structures and Related Concretions in Geochemical Studies
CIC crack fill and rims within concretions may have differing isotopic ratios. Based on δ13C values, Criss et
al. (1988) suggested that calcite CIC may form rapidly around decomposing organic matter, mainly
animal tissues and the resultant adipocere wax. Fish bones are also evident in those concretions
examined in that study. Criss et al. (1988) also concluded that early dolomitization of calcified adipocere
may have preserved initial marine environment temperatures, based on δ18O values.
As concretions undergo further alteration from infiltrating water and other fluids, these values will be
altered, but may still retain a traceable progression of that alteration back to their original δ13C and δ18O
values (Criss et al., 1988). This progression may also be time-correlative with vugs in adjacent rock.
Black-shale-hosted CIC showed δ18O values reflecting marine water and were coupled with δ13C values
indicative of fermentation and methanogenesis processes (Heimhofer et al., 2017). Chemistry in black
shale facies also indicated sulfate-reducing bacteria by the high abundance of pyrite.
Clumped isotope thermometry performed by Heimhofer et al. (2017) also suggested a formation depth
of only 650-850 meters, which is about half that expected from previously existing datasets. This
interpretation suggests that bacteria can facilitate calcite CIC formation from solutes in cooler marine
environments, and that CIC does not necessarily require deep burial diagenesis to provide heat and
Heindel et al. (2015) concluded that CIC REE and yttrium patterns indicate that calyx-shaped crystals in
their study precipitated from anoxic pore waters. Their study also concluded that benthic microbial mats
Figure 17: Reconstruction of the volcanic arc and potential eruptive source for North American Millbrig K-bentonite, as published
by Bergström et al. (2004)
in the sediment offered a hypersaline, low-oxygen marine environment in which CIC could form
As mentioned above, Hendry (2002) showed that calcite cements and BF-type CIC had isotopic, as well
as Fe+2 and Mg+2 values directly related to the adjacent shales. Hendry concluded that biologic activity
was not an initial factor in the formation of BF-type CIC based on δ13C values. Other studies have shown
abundant evidence that biologic activity probably is an initial factor in non-BF-type CIC. This helps
reaffirm that there are probably multiple methods to form CIC.
5.3. Serpent Mounds Impact Structures
It is worth quickly mentioning that the Serpent Mounds Impact Crater is just across the state boundary
from the collection site (Carlton et al., 1998). We briefly mentioned above that CIC structures are
optically related to impact cones. Many authors referenced herein have also noted that BF-type CIC are
commonly found in faulted areas, especially after release of pressure. It is therefore possible that the
nearby bolide impact at approximately 320 Ma is partially related to local pressure-induced dissolution,
and subsequent recrystallization. That research option is beyond the scope of this paper.
6. Other Bentonite Clay Settings
In this section, we offer some global examples of bentonite beds adjacent or related to CIC structures. In
many of these cases, CIC structures are also adjacent or related to other types of concretions, which
may also contain radial varieties of CIC.
6.1. Cretaceous Lower Barren Unit of the Graneros Shale, North Dakota
Contains 25 or more beds near Pueblo, each being a quarter of an inch to 4 inches in thickness. Some
are limonitic or contain jarosite. Some beds contain concretions composed of non-fossiliferous
amorphous limestone and calcite having CIC structure (Cobban & Scott, 1972).
6.2. Cretaceous Thatcher Limestone Member of Graneros Shale, North Dakota
Contains CIC structures at tops and bottoms of individual bentonite beds, stained with iron minerals.
The limestone is generally silty and unbedded. Limestone fractures conchoidally. Some 15 inches thick
(Cobban & Scott, 1972).
6.3. Cretaceous Pierre Shale, South Dakota
Contains bentonite beds with adjacent septarian concretion and CIC beds (Lavington, 1933; Gill et al.,
6.4. Devonian Jeffersonville Limestone, Indiana
A poorly indurated K-bentonite bed can be found as the Tioga Bentonite Bed (Doheny, Droste, and
6.5. Ordovician Colvin Mountain Sandstone, Alabama
Contains CIC structures and is adjacent to bentonite beds. The Deicke and Millbrig bentonite layers were
formed from a volcanic eruption during the Taconic orogeny during the Late Ordovician. These units are
members of the Decorah Shale, which infills the Decorah Structure, a possible impact structure with a
crater diameter of 6 km (French et al., 2018).
6.6. Late Devonian Yılanlı Formation, Turkey
K-bentonites are associated with diagenetically altered basaltic tephra in Turkey (Günal-Türkmenoğlu et
6.7. Archean Francevillian Basin, Gabon
Unmetamorphosed K-bentonite beds of calc-alkaline felsic volcanic origin show low-temperature S-I
conversion over long periods of time (Bankole et al., 2018).
Based primarily on the literature review, we make the following conclusions:
1. While CIC structures possibly form in many ways, more recent chemical analysis shows that CIC may
be primarily formed by the dissolution of smectite and neoformation of calcite and adjacent
smectite-illite shales. We propose that such a high calcium source must be near the CIC nucleation
2. We have summarized some strong evidence that the structure of CIC may be biologically driven, and
possibly related to sulfate-reducing bacterial mats hosting stromatolitic organisms.
3. Those summaries identified organic gels found within CIC and suggests the form of CIC may be
directly related to the radial growth pattern of trapped stromatolitic bacteria in a clay bed, either in
a peat bog, a shallow marine setting, or high pH lacustrine environment.
4. Our summary shows that the chemistry of CIC structures retains material that strongly suggest a
pyroclastic origin, including tuffaceous chemistry, and abnormally high quantities of radioactive
5. Studies have long shown that tonsteins are directly related to ash beds and have been used
extensively in the coal industry for stratigraphic correlation, and in the paleontological industry to
determine age of deposition and direction of winds in paleoenvironments. No evidence was found in
existing literature to directly relate tonsteins with marine CIC stuctures.
6. Abundant evidence exists to suggest that tonsteins and marine carbonate CIC structures are possibly
the same base material but evolved in differing environments, however, no evidence was found in
existing literature to directly show that tonsteins could host calcite CIC structures. Tonsteins may
instead contain relatively similar patterns formed from kaolinite crystals. This may simply be a
matter of pH during diagenesis.
7. Modern studies have shown via physical experimentation and clumped isotope thermometry that
CIC structures do not require deep burial or geologically high temperatures to form. Thermometry
showed that shallow burial and biologic activity did not even force isotope values to that of boiling
water. Instead, values suggested that biochemistry was able to complete diagenesis without
external heat or pressure while marine or high pH porewater was available in the substrate. Another
option given was low temperature hydrothermal input paired with bacterial activity (Farias et al.,
8. Soft sediment-hosted biologic activity is likely the cause of non-BF-type CIC structures, and allows
for easier displacive growth of both aragonite-calcite transitions, and smectite-calcite diagenesis, by
both organic and inorganic pathways.
9. This mechanism should function similarly in peat bogs, playa, and marine settings, though initial
mineral chemistry and organic activity will clearly play a part in determining diagenetic mineralogy.
Bogs are low pH, while playa environments are high pH. Stromatolitic bacteria, which are known
extremophiles, may offer an organism which could survive both environments.
Based on examination of our Kentucky sample, we make the following conclusions:
1. The sample shows a common CIC structure and traps Fe- and Zn-rich clays and sulfides in annual
rings. In addition, quartz-rich material is also trapped, which may have originally been volcanogenic
lapilli. Mg- and Ca-clays appear to be lacking or too small to identify, even at 400x magnification. We
conclude that if there were any to begin with, diagenesis of those clays spent their entire mass in
the conversion to carbonates.
2. The sample shows a high-K content in its clays based on ED-XRF data, which is normal for clays. In
addition, there is a remnant of Rb and Sr in the sample. Data suggests either that the sample
originally came from high-K feldspathic pyroclastic material, or that it accumulated abundant K in a
marine environment during transition from ash (or existing smectite) to illite. We conclude that the
original material was at least clay, if not volcanic ash.
3. ED-XRF and microscopy showed that pyritized organic matter acted as nucleation points for some of
the CIC structures. Other cones may also show this correlation in our sample, but the sample cut
does not allow for every cone to be investigated directly. This observation suggest that our sample’s
CIC structures may nucleate on biologic material as previously suggested in other studies.
4. ED-XRF and microscopy also showed that pyritized peloids are trapped adjacent to clays and quartz
in annual layers, allowing for internal biologic activity during formation, as evidenced by previous
Based on the marine deltaic environment of the Bedford Shale in Kentucky, and its proximity to the still-
forming Acadian range, there is a very large possibility that a feldspathic volcanic arc could be the source
of ash or ash-like material deposited within the delta. Based on other studies, specifically those
surrounding the Ordovician Millbrig K-bentonites, we conclude that if this is ash, then it should be
traceable to its source by the extent and thickness of the ash-fall zone.
Pindell and Dewey (1982) proposed such an arc to explain the Fire Clay tonsteins of Kentucky, which
formed during the late Carboniferous period, prior to the collision of the Amazonian Craton with the
North American Craton. It is possible that one or more such arc already existed in the Upper Devonian
and Lower Carboniferous periods. Ultraplinian eruptive points from the Ordovician period may also have
still been active in the latest Devonian.
7.1. Future Research
Further examination of the host rock in Kentucky, as well as in other adjacent states, should show
correlation in a similar fission to non-marine tonsteins. Based on literature review, we should be able to
track Ohio, Bedford and Sunbury shales over most of Ohio, and into Michigan.
To conclusively show that tonsteins are directly related to CIC beds in shale, or that CIC beds are related
to bentonite beds, we must find a location where marine shales grade into either peat bog tonsteins, or
where CIC grades directly into bentonite. The correlation between tonstein and bentonite was already
made by Senkayi et al. (1984).
We feel strongly that such a site must exist somewhere in North America, if not multiple locations.
7.2. Pennsylvanian CIC of Grand Ledge
A second sample obtained from float in Grand Ledge, Michigan shows the same pattern of
crystallization, but is far darker and denser than the Kentucky sample. A quick mention by Kelly (1936) in
a Michigan publication discusses the possible source of this float sample, and so we intend to investigate
that location in the spring of 2020.
If we can find the host rock, we can then determine if this is a tonstein or a marine bed. The host rock is
supposed to be Pennsylvanian in age, and is adjacent to coal, though the details of bed scale and order
are not clear. The hosting Saginaw Formation was determined to be a mixture of both marine shales and
continental peat bogs and sandstones at the time of that report, so may provide the exact transition
zone we’re looking for.
If successful, we may have evidence of a tonstein with CIC structure matching the physical
characteristics of marine shale-hosted CIC. The other option is that we identify a marine carbonate CIC
adjacent to a peat bog setting.
The sample itself may show the transition we’re looking for. Initial inspection of the Grand Ledge sample
shows spherical glassy beads within the clay packets between cones. Clays are located in the same
interstitial zones as in the Kentucky sample, and in the same scaled patterns, however, the Kentucky
sample did not retain clear spheroid shapes in its quartz grains. Initial ED-XRF data also shows a stronger
ferromagnesian signature than that of the Kentucky sample.
The Grand Ledge sample also shows a weathering rind, with a very clear Fe-weathering front. Below the
weathering front, glassy beads are better preserved, and interstitial clay appears to be intact and sealed.
Above the weathering front, glassy beads are sparse and loose, while clays are often missing, causing
the outside to become more brittle and crystalline in appearance.
Future tests on the Grand Ledge sample should include multiple geochronological examinations, and
XRD analysis to determine exact mineralogy of both the inner and outer chemistry. Isotopic analysis
should be performed on the Grand Ledge sample to determine if the weathering rind is indicative of the
of inner preservation, or if multiple stages of infiltration by groundwater are evident and contaminating.
The direct stratigraphic correlation of tonsteins, bentonite beds, and CIC structures in similarly thick
beds would solidify the idea that CIC is formed after volcanic ash. That ash, or its diagenetic remnants,
may be used in future studies as a proper geochronological bed. In addition, as has been done with
tonsteins, individual eruptions and their geographic source could be determined, with CIC bed thickness
added to the existing tonstein thickness data used in such studies.
In addition to stratigraphic correlation, CIC beds may show further evidence that feldspathic volcanic
arcs erupted continuously during the entire formation of Pangea. In addition, such eruptions may be
capable of producing the same formations today in playa and bog settings.
The implied and somewhat evidenced association with stromatolitic bacterial mats in CIC and BF
structures could give further evidence to the presence of life in Mars if similar structures are found
We thank Dr. Stephen Kaczmarek of Western Michigan University’s Geological and Environmental
Sciences Department (WMU) for the use of his lab’s thin section of the Kentucky sample. We also thank
Dr. Kaczmarek for allowing us to use his lab’s ED-XRF machine.
We also thank the Michigan Geological Repository for Research and Education (MGRRE) for the use of
their tools, time, and space. MGRRE Director and Professor Emeritus, Dr. William B. Harrison III has been
very informative on both Michigan and Kentucky geologic topics.
Thanks also goes to Brooks Ryan for assisting in obtaining the thin section slide, as well as previous
opportunities in learning to use the ED-XRF machine.
We thank Dr. Andrew Caruthers for pushing us toward taking on this project based on a single rock that
stood out in the mud.
We also thank Dr. Peter Voice for providing timely and related educational content, especially of which
breaks long-held belief systems in the geologic sciences. Many of the clay-related lectures were
instructional in the writing of this paper.
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