THE GEOCHEMICAL AND MINERALOGIC EXPRESSION OF SEQUENCE
BOUNDARIES IN MISSISSIPPIAN CARBONATES OF THE APPALACHIAN BASIN,
GEORGIA AND TENNESSEE
by
DANIEL EDWARD BULGER
(Under the Direction of PAUL SCHROEDER)
ABSTRACT
The clay mineral corrensite is a regularly ordered interstratified chlorite-smectite that is the result of low temperature diagenesis of Mg-rich minerals. Chemically, the smectite to chlorite transition through corrensite involves a decrease in alkali and alkaline earth metals with decreasing Si/(Si+Al), increase in Al+3 for Si+4 substituition in tetrahedral sites and variability in
Fe/(Fe+Al) of octahedral cations, which appears to be strongly influenced by whole-rock
composition. High resolution transmission electron microscopy has revealed two possible
mechanisms for the chloritization of smectite.
A comparison of two contrasting depositional sequences was conducted to test the
potential for corrensite as a proxy for sequence boundary identification in carbonate sequences
deposited in arid-semiarid marine environments. X-ray diffraction analysis of clay minerals
within the normal marine sequence of the Tuscumbia and Monteagle Formations reveal the
presence of corrensite in association with sequence boundaries and late highstand systems tract
(HST) deposits. Transgressive systems tract (TST) and early HST deposits representative of
normal marine conditions contain smectite, illite and minor kaolinite. In contrast, corrensite dominates the TST and HST of the restricted marine sequence of the Saint Louis Formation.
The Reelsville unit exposed at Jellico, Tennessee is a depositionally complex, well- developed unconformity bounded unit. Field and petrographic observations combined with clay mineral, stable oxygen and carbon isotopes and quantitative bulk rock mineral analysis was conducted on single hand samples. Excursion patterns in mean δ13C and δ18O values and dolomite concentration correlate with features indicative of subaerial exposure observed in the field. In addition, high variance in δ13C and δ18O values occurred at these positions with one exception. Six surfaces of subaerial exposure in the Reelsville unit were identified.
Petrographic observations of faunal diversity combined with traditional facies and mineral analysis suggest that the Reelsville unit was deposited in an inner ramp setting.
Petrographic analysis and field observation identified three parasequences in the Reelsville unit.
Diverse faunal elements suggest the base of parasequence occurred under open marine salinity conditions; however, massive dolomitized mudstone at parasequence tops suggests deposition under restricted conditions. The occurrence of corrensite throughout the Reelsville suggests deposition under high salinity conditions.
INDEX WORDS: Sequence boundary, Tuscumbia Formation, Monteagle Formation,
Reelsville unit, Dolomite, Carbon and oxygen isotopes
THE GEOCHEMICAL AND MINERALOGIC EXPRESSION OF SEQUENCE
BOUNDARIES IN MISSISSIPPIAN CARBONATES OF THE APPALACHIAN BASIN,
GEORGIA AND TENNESSEE
by
DANIEL EDWARD BULGER
BS, Bradley University, 1987
MS, Northeastern Illinois University, Chicago, 2002
A Dissertation Submitted to the Graduate Faculty of The University of Georgia in Partial
Fulfillment of the Requirements for the Degree
DOCTOR OF PHILOSOPHY
ATHENS, GEORGIA
2014
© 2014
DANIEL EDWARD BULGER
All Rights Reserved
THE GEOCHEMICAL AND MINERALOGIC EXPRESSION OF
SEQUENCE BOUNDARIES IN MISSISSIPPIAN CARBONATES OF THE APPALACHIAN
BASIN, GEORGIA AND TENNESSEE
by
DANIEL EDWARD BULGER
Major Professor: Paul Schroeder
Committee: Steve Holland R. Bruce Railsback Gerald Kuecher
Electronic Version Approved:
Maureen Grasso Dean of the Graduate School The University of Georgia May 2014
iv
ACKNOWLEDGEMENTS
I would like to thank all the people that made this project possible. First, I would like to
thank the members of my dissertation committee, Dr. Schroeder, Dr. Holland, Dr. Railsback and
Dr. Kuecher for their participation in this study as members of my committee. I am indebted to
Dr. Schroeder for serving as my advisor under difficult circumstances. In addition, if it weren’t
for Dr. Schroeder’s clay mineralogy class project, I would have never been introduced to the clay
mineral corrensite, which is the focus of this dissertation. I would like to thank Dr. Holland for his guidance throughout my graduate studies as well as introducing me to the concepts of
sequence stratigraphy, which defines the framework of this study. I thank Dr. Railsback for
introducing me to carbonate petrography and geochemistry. The knowledge I gained from his
classes has been invaluable. I appreciate Dr. Railsback open door policy and allowing me to
randomly show up in his office to talk over topics that clearly took away from his own work. I
would like to thank Dr. Kuecher for his knowledge and support over these years, which extends
back to my days at Northeastern Illinois University. It was Dr. Kuecher who first kindled my
interest in clay minerals. I would like to thank all the organizations that provided student
research grants for this project. This list of contributors includes the following: Gulf Coast
Association of Geological Societies, Society for Sedimentary Geology, Geological Society of
America, Society of Petrophysicists and Well Log Analysts, Clay Mineral Society, GDL
Foundation and Miriam Watts-Wheeler Graduate Studies Student Fund.
iv
TABLE OF CONTENTS
Page
ACKNOWLEDGEMENTS...... iv
LIST OF TABLES...... viii
LIST OF FIGURES ...... ix
CHAPTER
1 INTRODUCTION...... 1
Importance of Study...... 1
Regional Background and Sequence Stratigraphy...... 2
Purpose of Study...... 5
Questions, Hypotheses and Methods...... 6
Dissertation Format...... 10
2 CORRENSITE...... 12
Introduction...... 12
Corrensite Structure ...... 13
Precursor Minerals to Corrensite ...... 18
Temperature of Formation and Stability...... 20
Rock Fluid Effects...... 21
Inorganic Reactive Interactions ...... 24
Organic-Inorganic Reactive Interactions...... 28
Continuous and Discontinuous Transition...... 29
iv
Chemical Characterization of the Smectite-Corrensite-Chlorite Transition...... 34
HRTEM of the Smectite-Corrensite-Chlorite Transition ...... 44
Conclusions...... 48
3 SEQUENCE STRATIGRAPHIC DISTRIBUTION OF THE CLAY MINERAL
CORRENSITE IN TWO CONTRASTING MISSISSIPPIAN DEPOSITIONAL
ENVIRONMENTS...... 50
Introduction...... 50
Corrensite as a Geochemical Proxy ...... 52
Facies Description...... 55
Sequence Stratigraphy Key Definitions...... 55
Sequence Thickness and Accomodation...... 64
Geologic Setting of North America during the Mississippian Period ...... 64
Regional Stratigraphic Framework...... 68
Methods...... 73
Results...... 74
Disscussion ...... 82
Corrensite as a Proxy for Sequence Boundary Position ...... 84
Confirmation of Facies Interpretation...... 85
Conclusions...... 85
v
4 GEOCHEMICAL AND MINERALOGIC EXPRESSION OF SUBAERIAL
EXPOSURE IN MISSISSIPPIAN CARBONATES OF THE APPALACHIAN
BASIN: EVIDENCE FROM THE PINE MOUNTAIN OVERTHRUST BELT IN
NORTH CENTRAL TENNESSEE ENVIRONMENTS ...... 87
Introduction...... 87
Regional Background (Sequence Stratigraphy)...... 94
The Pine Mountain Outcrop...... 95
Physical Evidence for Lowstand Deposits and Subaerial Exposure...... 98
Structures Associated with Subaerial Exposure ...... 100
Methods...... 105
Results...... 112
Discussion...... 128
Other Surfaces of Subaerial Exposure ...... 137
Conclusions...... 140
5 SUMMARY OF RESULTS ...... 142
REFERENCES ...... 148
APPENDICES
A Microprobe analyses of expandable mafic phyllosilicate minerals ...... 181
B X-ray diffraction of bulk powder...... 201
C X-ray diffraction of clay mineral (<2µm) fraction ...... 292
D Summary of δ13C results from the upper C8/C9 sequences...... 382
iv
LIST OF TABLES
Page
Table 2.1: Ordered mixed-layered chlorite minerals (R1) and randomly interstratified chlorite
minerals (R0) ...... 14
Table 2.2: Corrensite temperature of formation and stability...... 22
Table 2.3: Summary of the chloritization process for biotite and smectite observed in HRTEM
studies ...... 47
Table 3.1: Carbonate and mixed carbonate-clastic ramp facies ...... 57
Table 4.1: Characteristics and stratigraphic position of seven physical features indicative of
subaerial exposure observed in the Reelsville unit...... 118
Table 4.2: Student’s t-test of difference of mean and f-test of difference of variance results of
δ13C and δ18O ...... 125
Table 4.3: Geochemical, physical and statistical evidence for the presence of subaerial exposure
found within the Reelsville unit...... 129
iv
LIST OF FIGURES
Page
Figure 1.1: Generalized structural elements of Kentucky and Tennessee...... 4
Figure 2.1: Differential thermal analysis curves...... 16
Figure 2.2: Inverse relationship between dolomite and chlorite content of Devonian lacustrine
mudrocks of the Orcadian Basin, Scotland...... 26
Figure 2.3: Relationship between vitrinite reflectance (% Ro), weight percent of chlorite (green
triangles) and dolomite (orange diamonds) in Devonian lacustrine mudrocks,
Orcadian Basin, Scotland...... 27
Figure 2.4: Variation of percentage of expandable layer in chlorite-smectite as a function of
depth from drillhole HT-42 of the Ohyu District, Akita Prefecture, Japan ...... 32
Figure 2.5: Mineral zonations of the basement complex La Palma seamount, Canary Islands ....33
Figure 2.6: Ternary plot of %Si, %Al and %Mg+%Fe+%Mn in mafic phyllosilicates
demonstrating a decrease in %Si and increase in %Al and %Mg+%Fe+%Mn during the
smectite to chlorite through corrensite transformation . All cation values are calculated
based on a chlorite-smectite formula with variable number of oxygens ...... 36
Figure 2.7: Plot of tetrahedral Al total vs Si/(Si+Al) in mafic phyllosilicates demonstrating a
strong negative correlation between these two variables...... 38
Figure 2.8: Plot of interlayer cation totals (Na+Ca+K) vs Si/(Si+Al) in mafic phyllosilicates
demonstrating a positive correlation between these two variables...... 39
Figure 2.9: Plot of Fe/(Fe+Mg) vs Si/(Si+Al) in mafic phyllosilicates demonstrating the lack of
correlation between these two variables ...... 41
iv
Figure 2.10: Plot of bulk rock Fe/(Fe+Mg) vs chlorite Fe/(Fe+Mg) for samples from Wales and
eastern North Greenland, demonstrating a strong positive correlation between these
two variables ...... 42
Figure 2.11: Plot of bulk rock Si/(Si+Al) vs chlorite Si/(Si+Al) for samples from Wales and
eastern North Greenland, demonstrating the lack of correlation between these two
variables ...... 43
Figure 2.12: Summary of three different chloritization mechanisms for biotite and smectite
observed in HRTEM studies...... 46
iv
Figure 3.1: Schematic representation of a corrensite superstructure subjected to different
treatments: A) Air dried; B) Ethylene glycol solvation...... 53
Figure 3.2: XRD patterns of the <2 µm fraction of a sample containing corrensite, illite and
minor kaolinite...... 54
Figure 3.3: Schematic facies profile for Mississippian carbonates of the Appalachian Basin...... 56
Figure 3.4: Photomicrographs of facies types for the Mississippian carbonates of the
Appalachian Basin ...... 58
Figure 3.5: Parasequence-stacking patterns in parasequence sets ...... 61
Figure 3.6: Type-1 and Type-2 sequence stratal patterns: Type-1 sequence and; Type-2 sequence
...... 62
Figure 3.7: Paleogeographic reconstruction; A) Late Mississippian time showing that the position
of Georgia and Tennessee within the dry trade-wind belt and; B) early Pennsylvanian
time illustrating the movement of the Appalachian Basin into the wet equatorial belt
...... 65
Figure 3.8: Foreland deformation model in which basin subsidence, peripheral upwarping and
forebulge migration (white arrows) are a response to thrust-belt loading (black arrow)
...... 67
Figure 3.9: Thickness of Mississippian sediments in the Appalachian Basin ...... 72
Figure 3.10: Sequence boundaries of the Lookout Mountain sequence ...... 75
Figure 3.11: Interpretive key for stratigraphic sections presented in Figure 3.12 and Figure 3.13
...... 77
Figure 3.12: Sequence stratigraphy and mineralogy of the Look Out Mountain locality in
Georgia...... 78
v
Figure 3.13: Sequence stratigraphy and mineralogy of the Pine Mountain locality in
Tennessee...... 80
Figure 4.1: XRD pattern for a regularly ordered (R=1) 50:50 interstratified trioctahedral chlorite and
trioctahedral smectite (corrensite) generated with the program Newmod: A) Air-dried
state; B) ethylene glycol-solvated ...... 93
Figure 4.2: Map of the outcrop locality on the Pine Cumberland overthrust sheet on I-75
southeast of Jellico, Campbell County, Tennessee...... 96
Figure 4.3: Stratigraphic column of the C8 sequence exposed on the Pine Cumberland overthrust
sheet on I-75 southeast of Jellico, Campbell County, Tennessee...... 97
Figure 4.4: The C8/C9 sequence boundary...... 99
Figure 4.5: Physical features of subaerial exposure...... 106
Figure 4.6: Representative example of the subsample collection scheme for each hand sample
...... 110
Figure 4.7: Facies interpretation of Parasequence 1 of the Reelsville unit (upper C8 sequence) at
the Pine Mountain locality...... 113
Figure 4.8: Facies interpretation of Parasequence 2 of the Reelsville unit (upper C8 sequence) at
the Pine Mountain locality...... 114
Figure 4.9: Facies interpretation of Parasequence 3 of the Reelsville unit (upper C8 sequence) at
the Pine Mountain locality...... 115
Figure 4.10: Physical evidence for subaerial exposure observed in the Reelsville unit...... 117
Figure 4.11: Bulk mineralogy, clay mineralogy, mean δ 13C and δ 18O, δ 13C and δ 18O variance
of samples collected from the Pine Mountain locality...... 120
Figure 4.12: Diffractogram of the clay mineral fraction (<2µm) from a sample interpreted as
normal marine (ooid grainstone) from the Beaver Bend Formation...... 122
vi
Figure 4.13: Statistical analysis of δ13C values...... 124
Figure 4.14: Statistical analysis of δ18O values ...... 127
Figure 4.15: Binary plots of variance vs sample area from samples of the upper C8 and lower C9
sequences from Pine Mountain locality: A) δ13C variance vs total sample area; B)
δ18O variance vs total sample area...... 136
1
CHAPTER 1 INTRODUCTION
Importance of Study
Sequence stratigraphy is a powerful methodology for predicting facies and thus the depositional control on the spatial distribution of petroleum reservoirs, seals and source rocks
(Posamentier et al., 1988). Most sequence stratigraphy studies have utilized lithofacies analysis; however, in recent studies involving siliciclastic sediments, the distribution of diagenetic alteration horizons have been useful in understanding the origin of spatial and temporal variations in the geochemistry of sediments (Al-Ramadan et al., 2005; Al-Ramadan et al., 2012;
2006; Ketzer et al., 2002; Morad et al., 2010; Salem et al., 2005; Taylor and Gawthorpe, 2003;
Taylor et al., 2004; Walz et al., 2012). This type of geochemical approach has enhanced our understanding of diagenetic processes within a sequence stratigraphic framework and, in turn, has created a high-resolution tool for the identification and interpretation of diagenetic processes within stratigraphically significant surfaces and systems tracts (Al-Ramadan et al., 2005; Al-
Ramadan et al., 2012; Burns et al., 2005; El-Ghali et al., 2009; El-ghali et al., 2006; Ketzer et al.,
2003; Ketzer and Morad, 2006; Ketzer et al., 2002; Masoumeh et al., 2011; Morad et al., 2010;
Salem et al., 2005; Walz et al., 2012). This research direction will ultimately help develop more comprehensive and precise, predictive models for the distribution of petroleum reservoirs and diagenetic seals (Ketzer et al., 2003; Morad et al., 2010). Despite advances in our understanding of these areas, many carbonate studies have focused primarily on the petrophysical aspects of 2 diagenetically altered carbonate (Ehrenberg et al., 2006; Fu et al., 2006; Hollis, 2011; Hopkins,
1999; Maliva et al., 2011; Petty, 2005; Ronchi et al., 2010; Vandeginste and John, 2013;
Vandeginste et al., 2009; Wierzbicki et al., 2006). Moreover, only a few studies have focused on the diagenesis of carbonate sediments within a sequence stratigraphic framework (Caron et al.,
2012; Csoma and Goldstein, 2012; Govert et al., 2012; Railsback et al., 2003; Railsback et al.,
2012; Saller et al., 1999; Schroeder et al., 2005; Smeester et al., 2012; Smeesters et al., 2003;
Taylor et al., 2000; Theiling et al., 2007; Westphal et al., 2004). In light of these issues, more work is needed to define the significance of diagenetic processes within a high resolution sequence stratigraphic framework over spatial and temporal scales.
Regional Background and Sequence Stratigraphy
Late Paleozoic tectonism in the Appalachian orogen is interpreted to be the result of the collision of North America with the combined African-South American continents (Scotese and
Mc Kerrow, 1990). The collision culminated in a series of large scale cratonward thrusting of
Paleozoic rocks onto the North American craton (Mack et al., 1983). The load of the thrusted sheets onto the craton produced the Appalachian foreland basin. The Cincinnati Arch and
Waverly Arch were topographic highs that were active during the Meramecian and Chesterian time (Dever, 1999; Ettensohn, 1980; Ettensoln, 1975; Pryor and Sable, 1974; Sable and Dever,
1990; Woodward, 1961). The Jessamine Dome in central Kentucky and the Nashville Dome in central Tennessee, situated along the long-axis of the Cincinnati Arch, are separated by a passageway between the Appalachian and Illinois Basins referred to as the Cumberland Saddle
(Figure 1.1). Faults, arches and domes active during the Mississippian resulted in rapid thickness changes across downdropped blocks and thinning-erosion over highs (Al-Tawil, 1997). 3
During the late Mississippian period, the northern edge of the research area (eastern
Kentucky) was situated at approximately 25 degrees south latitude (de Witt and McGrew, 1979;
Scotese and Mc Kerrow, 1990) within the dry trade-wind belt. By the early Pennsylvanian, the northward movement of the North American plate placed the basin in the wet equatorial belt (de
Witt and McGrew, 1979; Scotese and Mc Kerrow, 1990). This gradual shift from a semi-arid condition during the Meramecian and earliest Chesterian epochs to a wetter climate by late Early
Chesterian time is recorded in pedogenic features associated with subaerial exposure surfaces in northeastern Kentucky (Ettensohn et al., 1988). The early semi-arid climate was a major factor responsible for caliche formation within paleosols; however, by the Early Chesterian time, the climate had become more humid, and the latest formed caliches were partially destroyed by dissolution, creating a leached, clayey residual soil on top of earlier caliche soils (Ettensohn et al., 1988).
High-resolution sequence stratigraphy of the Mississippian succession in central
Kentucky has been determined by Al-Tawil and Read (2003) and provides the sequence framework for the Appalachian Basin. The Mississippian succession in eastern-central Kentucky is dominated by twelve small-scale fourth-order sequences, 5 to 20 m thick, bounded by relatively extensive disconformities. Regional disconformities are marked by paleosols, caliche, micro-karsting, brecciation, tepee formation or sharp contacts between limestone and overlying transgressive marine shales. The fourth-order sequences are stacked into several siliciclastic- bounded third-order or composite sequences that comprise broadly transgressive Mississippian supersequences. The sequences are relatively thin in the study area on the distal foreland and thicken significantly towards the east, the proximal foreland region in Virginia and West
4
Figure 1.1: Generalized structural elements of Kentucky and Tennessee.
5
Virginia, where they are 30 to 90 m thick (Al-Tawil et al., 2003). The sequences in the study area span 250 to 700 ky. Most of the high-frequency sequences can be correlated between the
Illinois Basin and the Appalachian Basin across the intervening Cincinnati Arch suggesting they reflect a combination of eustatic sea-level change and tectonic activity that appear to have controlled the timing of the sequences. The facies stacking in the lower sequences indicates they formed during moderate sea-level changes that effectively flooded the ramp to depths of 10 m or less (sequences 1 through 8). Facies stacking in the upper sequences suggests that magnitude of fourth-order sea-level changes increased in the later Chesterian, resulting from the flooding of the ramp to depths of tens of meters (sequences 9 through 12). This increase in magnitude was probably synchronous with the buildup of ice sheets on Gondwana by eccentricity forcing during a transition into ice-house conditions. Concomitantly, this change was accompanied by increasingly humid conditions, which decreased oolite deposition on the ramp and favored deposition of open marine skeletal limestones bounded by marine siliciclastic units.
Purpose of Study
It is the primary purpose of this dissertation to develop a diagenetic model that uses mineralogical and geochemical horizons indicative of sequence boundaries that could be used to help reconstruct basin histories. The primary objectives of this dissertation consist of the following: 1) Characterize geochemically the Mg-smectite to Mg-chlorite transition and describe the factors that affect the process; 2) identify diagenetic minerals that convey relevant information concerning the position of sequence boundaries; 3) determine whether heterogeneity in geochemical expression of stable carbon and oxygen isotopes exists at the scale of a hand- sized sample and trace the temporal variations in mineral and stable carbon and oxygen isotope 6 composition to assess climatic conditions during deposition. During the development of the diagenetic model the following questions will be answered by testing the following hypotheses:
Questions, Hypotheses and Methods
Question 1:
What factors affect the smectite-corrensite-chlorite transition and does the transition differ geochemically in hydrothermally altered mafic igneous rocks compared to thermally altered sedimentary rocks.
Hypothesis 1.1:
The smectite-corrensite-chlorite transition will proceed geochemically in a similar manner in different diagenetic environments.
Methods 1.1:
A review of the literature will be conducted to determine the factors that affect the smectite-corrensite-chlorite transition such as: 1) temperature of formation; 2) rock-fluid ratio; 3) redox environment and; 4) host rock geochemistry. Electron microprobe data reported in the literature are used to assess changes in composition during the smectite-to-chlorite transition.
Two contrasting diagenetic environments are selected: hydrothermally altered igneous/volcanoclastic and thermally altered sedimentary.
Question 2:
Do the physical expressions of diagenetic processes provide useful information in sequence stratigraphy studies? Current methodologies for identifying surfaces of subaerial 7 exposure in carbonate rocks utilize replicate sampling and analysis of δ13C and δ18O values and
Sr trace element concentrations (Theiling et al., 2007). However, can clay mineral suites
analyzed with XRD from single samples convey the same information concerning the position of
sequence boundaries?
Hypothesis 2.1:
Corrensite can be used to identify the position of subaerial exposure and peritidal facies
associated with periods of arid-semiarid climate when Mg-rich precursor minerals are present.
Methods 2.1:
Samples will be collected along a vertical transect at approximately 30 centimeter
resolution from two contrasting depositional sequences: 1) a restricted marine sequence
deposited under limited accommodation space and; 2) a normal marine sequence deposited under
relatively greater accommodation space. The M1 sequence of the Saint Louis Formation at the
Pine Mountain outcrop of Al-Tawil and Read (2003) represents the restricted marine sequence.
A sequence exposed in a 12 m section in an abandoned quarry on the southeast flank of Lookout
Mountain in Walker County, Georgia serves as the normal marine sequence. Clay minerals were
separated from carbonate rock with a buffered acetic acid solution. The <2µm fraction was by
centrifugation. An oriented clay mineral slide was prepared by pipetting the clay slurry onto a petrographic slide and allowing the slide to dry. The slides were analyzed for clay mineralogy with a Bruker D8 Advance XRD using the following parameters: 2 to 35o 2 Θ, 40 ma, 40 kv at 5o
per minute step and 0.02 increment. Clay mineral slides were run both dry and ethylene glycol saturated. Regions where corrensite is detected was mapped on a stratigraphic section diagram.
The occurrence of corrensite will then be compared with the position of previously defined 8 sequence boundaries. These results will define the relationship between the occurrence of corrensite and sequence boundaries.
Hypothesis 2.2
Quartz, calcite and dolomite percentages can be used to identify the position of subaerial exposure surfaces at the Pine Mountain locality. The quartz percentage will be greater near exposure surfaces due to the greater influx of siliciclastics from nearby weathering sources, reduced production of carbonate in restricted marine settings from biogenic sources and precipitation of diagenetic microcrystalline quartz from the dissolution of metastable siliciclastic sediments. Dolomite percentage will be greatest in restricted marine settings below sequence boundaries. Calcite percentage will be greatest in normal marine facies where calcite producing biogenic sources dominate.
Methods 2.2
The presence of quartz, calcite and dolomite will be quantitatively assessed with XRD.
Bulk sample will be lightly crushed. Approximately 8 g of the less than 2mm fraction was placed in a McCrone Mill with ethanol and ground for 10 minutes. Randomly oriented bulk powders were analyzed with a Bruker D8 Advance XRD under the following parameters: 2 to 75o 2 Θ, 40 ma, 40 kv at 5o per min step and 0.02 increment. Percentages for quartz, calcite and dolomite
were calculated with a Rietveld refinement algorithm provided in the program Topas ® by
Bruker. Quartz, calcite and dolomite percentages were compared with the position of sequence
boundaries, depositional environments identified by Al-tawil and Read (2003) and clay mineral
suites identified with XRD.
9
Question 3
Is sample variance for stable C and O isotopes in hand-sized samples great enough to recognize surfaces of subaerial exposure (sequence boundaries) in Mississippian age carbonate rocks carbonate?
Hypothesis 3.1:
Increases in variance in samples below exposure surfaces for stable C and O isotopes can be recognized in relatively small hand-sized samples. Theiling et al. (2007) recognized that variance for stable C and O isotopes increased in carbonate samples collected below surfaces of subaerial exposure. This phenomenon was attributed to permeability variability of the substrate below the exposure surface as well as variability in the distribution of photosynthesizers colonizing the surface above. The sampling scheme of Theiling et al. (2007) involved collecting replicates two meters apart along each horizon (a total of 10 samples per horizon); however, variance within a single hand sample was not tested.
Methods 3.1
Detailed, high resolution stable carbon and oxygen isotope concentration profiles were made for a sequence boundary that displays complex facies relationships and multiple surfaces of subaerial exposure below the boundary (Reelsville unit associated with the C8/C9 sequence boundary of Al-tawil and Read, 2007). Vertical transects were spaced approximately 30 cm from each other horizontally. Fist-sized samples will be collected perpendicular to the outcrop face and slabbed with a water cooled tile saw. The slab surface was gridded into ten 1cm2 areas. A
dental drill was used to isolate micrite from each cell of the gridded region on each slab. Drill
speeds were be kept low to avoid alteration effects related to high speed drilling (Gill et al., 10
1995). Variance in δ13C and δ18O values for the ten subsamples were calculated and compared
with the position of subaerial exposure surfaces identified by Al-Tawil and Read (2003) and
statisitically significant excursions in δ13C and δ18O described in Methods 1.1 section. The δ13C
and δ18O values from each sample were evaluated statistically so that the position of the C8/C9
sequence boundary and associated surfaces of subaerial exposure identified during field
observation could be confirmed geochemically. Surfaces of subaerial exposure within the current
sequence stratigraphic framework and surfaces previously unidentified were placed in context
with clay mineral suites identified by XRD such that δ13C and δ18O value excursion patterns could be used to evaluate the conditions that influence the occurrence of these diagenetic horizons.
Dissertation Format
Results of this study have been organized as three separate manuscripts. Chapter Two addresses specifically the geochemical transformation aspect of the Mg-smectite to Mg-chlorite transition. Geochemical analyses from ten different studies are used to compare and contrast the chemistry of the smectite-corrensite-chlorite transition between altered mafic igneous and sedimentary rocks. This chapter specifically addresses Question 1 presented in Questions and
Hypotheses section of this chapter. Chapter Three links clay mineral diagenesis to a high resolution sequence stratigraphic framework and illustrates how this approach can provide valuable information concerning the position of sequence boundaries deposited within semiarid- arid settings. This chapter addresses Question 2 presented in the Questions and Hypotheses section of this chapter. Chapter Four specifically tests the idea that variance in stable carbon and oxygen isotopes in hand-sized samples can be used to identify the position of sequence boundaries. In addition, calcite and dolomite concentrations and corrensite are used to infer 11 environmental conditions at the time of deposition of a well-developed, complex sequence boundary. Chapter 4 addresses Question 3 presented in the Questions and Hypothesis section of this chapter. Chapter Five synthesizes information in previous chapters and summarizes the major conclusions of the study.
12
CHAPTER 2
CORRENSITE
Introduction
Corrensite was first described by Stephen and MacEwan (1951) as “swelling chlorite” only later to be renamed “corrensite “ by Lippman (1954) in honor of his advisor C.W. Correns
(Moore and Reynolds, 1997). Corrensite occurrence has been described in a variety of geological contexts including thermally and hydrothermally altered volcanoclastic sediments and igneous rocks (Beaufort et al., 1997; Bettison and Schiffman, 1988; Bevins, 1991; Inoue and
Utada, 1991; Inoue, 1984; Jiang and Pecor, 1994; Jimenez-Millan, 2008; Kameda, 2011; Kubler,
1973; Mas, 2008; Morrison and Parry, 1986; Murakami et al., 1999; Robinson and Santana de
Zamora, 1999; Schiffman and Staudigel, 1995; Son et al., 2001; Stephen and MacEwan, 1951), ancient marine carbonates, siliciclastics and evaporites (Andreason, 1992; Bayhan, 2007;
Bodine, 1985; Bradley and Weaver, 1956; Drits, 2011; Droste, 1963; Eardley et al., 1956; Fisher,
1988; Han et al., 2000; Hillier, 1993; Hints et al., 2006; Jeans, 2006; Jeans et al., 2005; Kopp and
Fallis, 1974; Monnier, 1982; Pay, 2000; Peterson, 1961; Rao, 1992; Ryan and Hillier, 2002;
Sandler, 2001; Stein et al., 1991; Trindade et al., 2013), ancient evaporative lacustrine settings
(April, 1981; Barrenechea, 2000; Bristow, 2009; Hillier, 1996; Lippmann, 1956; Sandler, 2001;
Yemane et al., 1996) and along subduction thrust/lateral fault margins (Kameda, 2011;
Schleicher, 2012; Schleicher et al., 2010). Since low charge corrensite, a regularly interstratified chlorite-smectite, is not a primary constituent of rock forming materials at the earth’s surface, its occurrence is believed to be the result of low temperature mesodiagenesis of a (Mg-rich) 13 precursor mineral. In addition, corrensite appears to represents a distinct phase thermodynamically in the prograde sequence smectite>corrensite>chlorite (Beaufort et al., 1997;
Drits, 2011; Inoue and Utada, 1991; Reynolds, 1988b; Shau et al., 1990; Son et al., 2001).
However, the details of the transition have been a topic of vigorous debate (Beaufort et al., 1997;
Bettison and Schiffman, 1988; Bevins, 1991; Chang et al., 1986; Inoue and Utada, 1991; Inoue,
1984; Kogure et al., 2013; Murakami et al., 1999; Reynolds, 1988b; Robinson and Santana de
Zamora, 1999; Sandler, 2001; Schiffman and Staudigel, 1995).
Elucidation of corrensite structure, chemistry and paragenesis is important because it offers insight into mineral diagenesis. The purpose of this chapter is to present an overview of the smectite-corrensite-chlorite paragenetic sequence with a general treatment leading to greater understanding of the occurrence of corrensite and its association with sequence boundaries in carbonate rocks deposited in arid-semiarid settings. Factors that affect the smectite-corrensite- chlorite transition include: 1) precursor minerals; 2) temperature; 3) rock/fluid ratio; 4) inorganic reactive interactions; 5) organic-inorganic reactive interactions and; 6) host rock geochemistry.
Electron microprobe data can be used to assess changes in geochemistry during the smectite- corrensite-chlorite transition. Two contrasting diagenetic environments will be discussed: 1) hydrothermally altered basalt and diorite, and 2) thermally altered sedimentary sequences.
Corrensite Structure
Two-component mixed layered clay minerals are defined by their phyllosilicate layer types and the ordering of those layer types. Chlorite forms two-component interstratified minerals (Table 2.1) with all of the other common phyllosilicate layer types (Reynolds, 1988b).
The term corrensite was originally proposed by Lippman (1954) to describes a regularly ordered
14
Table 2.1: Ordered mixed-layered (R1) and randomly interstratified chlorite minerals (R0).
15
(50:50) interstratified chlorite and swelling chlorite clay mineral. The 14Å spacing of the chlorite layer adjacent to a 12.5-17Å swelling chlorite layer results in a superstructure that is 31Å thick when fully expand which is easily identifiable on X-ray diffractograms. However, there is some variability in the definition in that the 50:50 interstratification of trioctahedral chlorite with either trioctahedral smectite or trioctahedral vermiculite are both considered “corrensite” by the
Nomenclature Committee of The Clay Minerals Society (Bailey, 1982; Guggenheim et al.,
2006).
The thermal behavior of corrensite has been widely investigated (Bradley and Weaver,
1956; de Kimpe et al., 1987; Lackschewitz et al., 2000; Lippmann, 1956; Peterson, 1961) and offers insight into the layer types found in this mineral. The differential thermal curve for high charge corrensite is a composite of the individual curves for vermiculite and chlorite (Figure
2.1A). Analyses of high charge corrensite reveal endothermic peaks at 500-600o C and 810-950
oC characteristic of chlorite and an endothermic peak at 200 oC characteristic of vermiculite
(Bradley and Weaver, 1956). Similarly, the differential thermal curve for low charge corrensite, a 50:50 interlayering of chlorite and smectite, is a composite of the curves of chlorite and smectite (Figure 2.1B). Endothermic peaks at 80o and 160o are typical of smectite in which
exchange sites are sarturated with divalent cations (de Kimpe et al., 1987).
The ordering of chlorite and smectite layers in the stacking sequence of corrensite is
described as having a Reichweite of 1 or “R1” as defined by Jadgodzinski (1949). R1 is a
statistical designation related to “reach-back”: the correlation distance over which the occurrence
of one layer type affects the probability of occurrence of another layer (Reynolds, 1988b).
Statistically, R1 means that only adjacent layer type affects the probability of an adjacent layer
resulting in an alternating layer-type structure. Considering corrensite as an example, if “A” 16
Figure 2.1: Differential thermal analysis curves: A) High charge corrensite (de Kimpe et al., 1987); B) low charge corrensite (Bradley and Weaver, 1987). The curve for both corrensite types is a composite of the curves of their respective interlayers.
17 represents the chlorite layer and “B” represents the smectite layer, the expected stacking pattern for an R1 chlorite-smectite would be ABABABAB…etc. A crystal composed of an ordered structure at 50:50 is not a mixed-layered mineral, and its fixed stoichiometry and structure qualify it as a distinct mineral species with a unit cell defined by the chlorite-smectite super structure (Moore and Reynolds, 1997; Reynolds, 1988a). In contrast, the designation “R0” means that adjacent layers have no effect on the occurrence of a given layer type and results in a random interstratification of the two components. Randomly interstratified (R0) chlorite-smectite minerals are considered mixed-layered clays and are identified by the mineral name and their relative proportions. For example, chlorite (0.6)-smectite designates a mixed-layered chlorite- smectite that is 60% chlorite and 40% smectite.
Regularly interstratified (R1) chlorite-smectite clay minerals can be distinguished from randomly interstratified (R0) chlorite-smectite clay minerals by X-ray diffraction (XRD). The diffraction characteristics of corrensite are the result of the summation of the 14Å chlorite layer with the neighboring expandable 12.5-19 Å smectite layer. The XRD features of corrensite can be summarized as follows: 1) air dried material gives rise to a (001) reflection between 26.5-31
Å, and 2) upon ethylene glycol solvation, the (001) reflection expands to 31-33 Å due to the swelling characteristics of the smectite layers. In contrast, randomly interstratified (R0) chlorite- smectite produces a (001) peak upon ethylene glycol solvation variably positioned between 14 and 16.9 Å depending upon the proportions of chlorite to smectite. XRD patterns from mixed- layer clay structures differ from R1 superstructures and relatively pure clay mineral samples by the presence of nonrational series of basal reflections. The basal node position within the nonrational series can be predicted using the model presented by Méring (1949) and is stated in the following linear relationship: 18
WA= y/(x+y)
where WA is the layer proportions of clay type A, and x and y are the distances between nodes
corresponding to the mixed layered structure and the nearest nodes corresponding to two
periodic layer components (Méring, 1949). The shape and position of the diffraction maxima
vary depending upon the proportions and positions of rational reflections from each of the two
components involved. A higher proportion of a particular interstratified mineral would shift the
maxima towards that minerals discrete reflection position and the width would depend upon the
distance between the rational peaks of the two separate components. As such, random
interstratification of component layers change the position and widths of the diffraction maxima,
which is a diagnositic characteristic of the R0 mixed-layered system.
Precursor Minerals to Corrensite
The occurrence of corrensite is the result of mesodiagenesis of Mg-rich precursor
minerals. Precursor minerals to corrensite may include Mg-rich smectites like saponite or
stevensite or involve reactions between Mg-rich carbonate (dolomite) and dioctahedral clay
minerals. Several studies of modern evaporative environments suggest that Mg-rich smectite
forms authigenically. The composition of smectite minerals can become modified by interactions
with brines in modern hypersaline playa environments (Ataman and Baysal, 1978; Hover et al.,
1999; Jones and Weir, 1983; Martini et al., 2002; Singer and Stoffers, 1980). For example, Mg-
rich smectite is found locally in Lake Chad basin, located in Northern Central Africa. The Mg-
smectite neoformation process is believed to be a major control of the magnesium concentration
in such lake systems (Tardy et al., 1974). Indeed, experimental evaporation of diluted water from 19 the Chari River, one of two rivers feeding into Lake Chad, results in the co-precipitation of a variable suit of amorphous magnesian silicates (Gac et al., 1977). It is further suggested that these amorphous Mg-silicates undergo neoformation to Mg-rich smectite.
In a study of a marine evaporative environment in Baja California (Hover et al., 1999), authigenic Mg-rich trioctahedral smectite (Saponite) formed from detrital aluminous dioctahedral smectite in pore waters having a relatively high Mg+2/Ca+2 ratio. These high ratios
are often achieved subsequent to gypsum precipitation in the absence of dissolved inorganic carbon (Hover et al., 1999; Martini et al., 2002), which allows for Mg+2 to be incorporated in
octahedral sites during a precipitation of a discrete authigenic phase (Mg-rich trioctahedral
smectite). In the presence of dissolved inorganic carbon, sulfate reduction proceeds and the
thermodynamic constraints to dolomite precipitation are minimized resulting in a lowering of the
Mg+2/Ca+2 ratio in pore water as dolomitization of pre-existing calcite progresses.
In ancient evaporative marine and fluvial settings, precursor mineral formation may be initiated the precipitation of authigenic trioctahedral Mg-rich smectite (April, 1981; Hillier,
1994, 1996; Hover et al., 1999; Jeans, 2006; Ryan and Hillier, 2002). April (1981) found
corrensite in Jurrassic lacustrine facies and argued that under relatively high pH Mg-rich fluids
can facilitate the precipitation of brucite-like layers within a precursor trioctahedral smectite.
These data are consistent with laboratory results of Slaughter and Milne (1960) who precipitated
brucite in the smectite interlayer in suspensions of montmorillinite in 1 to 2 N MgCl2 with LiOH.
A high-resolution study of clay mineralogy, textures, structures, and compositions in sediments
from a transect across an arid Plio-Pleistocene rift basin, Olduvai Gorge Tanzania, showed that
smectite-rich clay minerals record geochemical signatures of fluvial and lacustrine depositional
environments (Hover and Ashley, 2003). For example, sediments from distal fluvial 20 environments, which were exposed to pedogenic processes, contained dioctahedral Al-rich smectite, whereas sediments from the lake margin and basin contained abundant Mg-rich smectite, which was an alteration or neoformation product from pedogenic dioctahedral Al-rich smectite. Jones and Weir (1983), Jones (1986) and Webster and Jones (1994) suggested that the precipitation of authigenic Mg-rich smectite in saline alkaline lakes is facilitated by precursor dioctahedral smectite that is acting as a substrate for neoformation. Results from Hillier (1994) and Hillier et al. (1996) suggested that pore-lining, Mg rich chlorites in aeolian/sabkha sandstones originated from precursor Mg-rich smectites formed from evaporative brines at near- surface conditions. Similarly, Ryan and Hillier (2002) stated that saponite, which commonly forms in aeolian and evaporative environments, was a precursor to corrensite, found in the uppermost sequence of the Sundance Formation which mark the regression of the Sundance Sea.
Precursor phyllosilicates to corrensite are ubiquitous components of volcanoclastic and intermediate-mafic igneous rocks subjected to low grade metamorphism. Trioctahedral smectite neoformation results from the low grade metamorphic alteration of these rocks. The occurrence of corrensite and chlorite in igneous rocks is believed to be the result of the neoformation of trioctahedral smectite associated with hydrothermal systems (Beaufort et al., 1997; Bettison and
Schiffman, 1988; Jimenez-Millan, 2008; Lackschewitz et al., 2000; Merriman, 2006; Robinson and Santana de Zamora, 1999; Schiffman and Staudigel, 1995), low grade metamorphism of volcanoclastic rocks (Inoue and Utada, 1991; Inoue, 1984; Murakami et al., 1999) and low grade metamorphism of intrusive igneous rocks and metabasalts (Bevins, 1991; Sandler, 2001).
Temperature of Formation and Stability
Literature describing the saponite-corrensite-chlorite transformation in relation to paleothermometry fall within two lithologic associations: 1) Metavolcanoclastic-igneous (Bylina, 21
2006; Inoue and Utada, 1991; Schiffman et al., 1991; Schmidt and Robinson, 1997; Versh et al.,
2005) and; 2) siliciclastic rocks (Chang et al., 1986; Hillier, 1993; Jeans et al., 2005; Miller and
Macdonald, 2004; Ryan and Hillier, 2002; Son et al., 2001; Stalder, 1979). While these settings vary, they are generally considered to be low-grade metamorphic environments that experienced hydrothermal alteration (Bylina, 2006; Lackschewitz et al., 2000; Schiffman et al., 1991;
Schmidt and Robinson, 1997; Versh et al., 2005), thermal alteration related to igneous intrusion
(Inoue and Utada, 1991), or alterations associated with a thermal gradient established during burial (Chang et al., 1986; Hillier, 1993; Jeans et al., 2005; Miller and Macdonald, 2004; Ryan and Hillier, 2002; Son et al., 2001; Stalder, 1979), the later of which contains a gentle gradient.
Methods for constraining the paleothermometry of saponite-corrensite-chlorite paragenic sequences include vitrinite reflectance (Chang et al., 1986; Hillier, 1993; Miller and Macdonald,
2004; Stalder, 1979), fluid inclusion microthermometry (Bylina, 2006; Versh et al., 2005), crystallization temperatures of indicator minerals (Bylina, 2006; Versh et al., 2005), oxygen isotope analysis of authigenic clays (Lackschewitz et al., 2000), Rock-Eval pyrolysis (Son et al.,
2001) and spore color index (Jeans et al., 2005). According to data summarized from the literature, a range of stability for corrensite exists between 60 and 300oC for a variety of
lithologic and diagenetic settings (Table 2.2).
Rock Fluid Effects
Factors such as rock permeability and water/rock ratios play an important role in the kinetics of
smectite-chlorite transformation sequence for hydrothermally altered mafic rocks and
sedimentary rocks (Bettison-Varga and MacKinnon, 1997; Chang et al., 1986; Chudaev, 1978;
Kogure et al., 2013; Miron et al., 2012; Shau and Peacor, 1992). Chudaev (1978) recognized that
the porosity of hydrothermally influenced flyschoid sediments of eastern Kamchatka had an 22
22
thermometry methodology, lithologic association and diagenetic and diagenetic association lithologic methodology, thermometry
and stability, formation temperature of Table 2.2: Corrensite described in the literature. environment 23 effect on the clay mineral composition. In this study, higher permeable sandstones primarily contained chlorite while lower permeable argillites contained either a corrensite-like mineral or montmorillonite and a random chlorite-smectite mixed-layer mineral. Chudaev (1978) concluded that clays formed in sandstone were structurally more complete and thermodynamically more stable than those formed in argillite. Chang et al. (1986) recognized a lower temperature regime for chlorite-smectite (C/S) ordering in sandstone compared to shales of Cretaceous sediments located in the northern continental shelf of Brazil that he attributed to the greater permeability of sandstone. Data from Shau and Peacor (1992) indicate that permeability controlled the degree of hydrothermal alteration of basalt at the DSDP Hole 504B located south of the Costa Rica
Rift in the Pacific Ocean, where well crystallized discrete trioctahedral phyllosilicate phases occurred within high-permeable veins whereas regions of low permeability contained mixed-layer clays with poorly defined packets of partially incoherent layers.
Bettison-Varga and Mackinnon (1988) performed a comparative study of vesicular and groundmass phyllosilicates in a hydrothermally altered basalt from the Point Sal ophiolite, and showed that randomly interlayered chlorite-smectite is a metastable phase formed in microenvironments of low water/rock ratio, while chlorite formed in microenvironments dominated by high water/rock ratio. Schmidt and Robinson (1997) studied the Keweenawan metabasaltic, North Shore Volcanic Group in Minnesota and concluded that areas of higher porosity and permeability are corrensite and chlorite dominant. In contrast, smectite is the dominant phase in massive metabasalts where flow centers are characterized by low porosity and permeability. Data from Schiffman and
Staudigel (1995) suggest that the pervasive alteration of extrusive and volcanoclastic 24 rocks of the La Palma seamount were evidence for high water/rock ratios and further suggest that the discontinuous smectite to chlorite neoformation pathway is favored under these conditions. Miron (2012) attributed the high variability of chlorite content in mixed-layered chlorite-smectite minerals found in mafic island-arc volcanics, south
Apuseni Mountains, Romania to slow reaction rates associated with low fluid-rock ratio.
Inorganic Reactive Interactions
Burial diagenetic reactions between clay minerals and other mineral components in sedimentary rocks have been used to explain the presence of chlorite. Although much of the work in the literature does not describe the presence of corrensite, the inorganic reactive interaction is presented here as a possible mechanism for the presence of corrensite. Zen (1959) originally proposed the following reaction to explain chemical equilibrium of minerals observed in modern sediments off the coast of Peru and Chile: dolomite + kaolinite + quartz + water = calcite + chlorite + carbon dioxide. This reaction was later used to explain the calcite-chlorite assemblage found metamorphic rocks of
Castleton, Vermont, (Zen, 1960). Muffler and White (1969) documented the replacement of dolomite, ankerite and kaolinite in sediments at shallow depth in the Pliocene and
Quaternary siltstones and sandstones of the Salton Sea geothermal field with an assemblage of chlorite and calcite at greater depth. They proposed the following reaction to explain the change in mineralogy: dolomite + ankerite + kaolinite + ferric iron + water
+ oxide = chlorite + calcite + carbon dioxide. Iijima and Matsumoto (Iijima, 1982) suggest that reactions between kaolinite and siderite during early diagenesis produce bertheirine which subsequently transforms into chamosite (Fe-rich chlorite) at elevated temperatures during burial diagenesis. Hucheon (1990) postulated a similar reaction 25 between dolomite and illite to explain mineral assemblages in Cretaceous sandstones of the Kootenay formation near Calgary, Canada, and in sandstones from the Venture Field of the Scotian Shelf, Nova Scotia. Hillier (1993) ascribe the inverse relationship between dolomite and chlorite content in mudrocks from the Devonian Orcadian Basin, Scotland, to burial diagenetic reactions between dioctahedral clay minerals and Mg-rich carbonates
(Figure 2.2 ). Further he stated that the inverse relationship is correlated with organic maturity: In many low maturity samples, dolomite is abundant and chlorite is absent; whereas at high maturity, chlorite is abundant and dolomite is minor or absent (Figure
2.3). Barrenchea et al. (2000) recognized a similar trend between chlorite and dolomite in the late Jurassic fluvio-lacustrine sediments of the Cameros Basin of northeastern
Spain. They concluded that chlorite and chlorite-mica crystallization involves a concomitant destabilization of dolomite, which is believed to be the Mg source for the neoformation of chlorite and chlorite–mica with little Fe+2. However, in a study of the
origin of clay mineral assemblages in the Germanic facies of the English Trias Jeans
(2005), found no evidence of mixed-layer smectite-chlorite minerals and Mg-rich chlorite
developing from metamorphic or diagenetic reactions between dolomite and detrital or
early diagenetic clay minerals. He concluded that mineral-lithofacies relationships are
best explained by competitive inhibition with carbonate and silicate which compete for
available Mg+2 during eodiagenesis. Indeed, results from Hover et al. (1999) and Martini
et al. (2002) illustrate how Mg+2/Ca+2 ratios in pore water can be suppressed during
dolomitization of pre-existing carbonate. This argument assumes that vertical and lateral
zonations within clay mineral assemblage are controlled by the physio-chemical setting
of the environment of deposition and intrinsic diagenesis (Jeans et al., 2005). 26
30
25
20
15
Dolomite% 10
5
0 0 5 10 15 20 25 30
% Chlorite
Figure 2.2: Inverse relationship between dolomite and chlorite content of Devonian lacustrine mudrocks of the Orcadian Basin, Scotland (from Hillier, 1993).
27
8
7
6
5 % Ro
4
3
2
1 0 5 10 15 20 25 30 Percent
Figure 2.3: Relationship between vitrinite reflectance (% Ro), weight percent of chlorite (green triangles) and dolomite (orange diamonds) in Devonian lacustrine mudrocks, Orcadian Basin, Scotland. An inverse relationship between chlorite and dolomite is related to organic maturity suggesting a reaction between dolomite and phyllosilicate minerals with increasing thermal regime (from Hillier, 1993).
28
Organic-Inorganic Reactive Interactions
The smectite-corrensite-chlorite series is also influenced by organic matter in rocks and subsequent decarboxylation during mesodiagenesis. Curtis (1978) examined the link between the organic and inorganic components separately during depth dependent geochemical reactions and the reduction of Fe+3 in siliciclastic rocks. Surdam
and Crossey (1985) suggest that f(O2) and f(CO2) conditions within some diagenetic
environments are controlled by oxidation-reduction reactions related to clay mineral neoformation and concomitant decarboxylation of kerogen. Under these conditions the reduction of ferric iron released during clay diagenesis, or within the clay structure itself,
is balanced by the oxidation of kerogen. Oxidation of kerogen results in the formation of
organic acids which further enhance the mobility of iron and aluminum in fluids near the
vicinity of the reactions. Indeed, experimental hydrothermal work (2 kbar water pressure,
300-400oC) by Velde (1977) on hydrothermally treated natural clay minerals show that
the stability of Fe-rich dioctahedral phyllosilicate minerals may be controlled by the
organic content of the starting material. He concluded that iron reduction was enhanced
by the presence of organic matter in rock, and this facilitated the precipitation of chlorite.
In contrast, poorly crystallized chlorite precipitated in the absence of organic matter.
Meunier et al. (1988) attributed differences between two alteration events in the “Les
Crêtes” Paleozoic granite in Vosges, France to changes in redox conditions where
corrensite and hematite was the stable phase under oxidative conditions, whereas chlorite
is the stable phase under more reduced (Fe+3 free) conditions. Results from Jiang and
Peacor (1994) suggest that the paragenetic relations observed in a prograde sequence of
pelitic rocks from the Gaspé Peninsula in Quebec were controlled by local variations in 29
f(O2) and f(CO2) governed by the decarboxylation of organic matter. Specifically, an
oxidized corrensite-magnetite-titanite assemblage was reduced to a higher grade chlorite-
pyrite-calcite-rutile assemblage under reducing conditions.
In a study of the fluvio-lacustrine sediments of the Cameros Basin of northeastern
Spain, Barrenechea et al. (2000) showed that corrensite and chlorite abundance were
related to redox conditions, where corrensite was favored in an oxidizing environment
and chlorite plus corrensite was favored under more reducing conditions. The
geochemistry of chlorites in this study shows a higher Fe+2 content towards the center of
the stratigraphic profile where organic matter content and the potential for a reducing
diagenetic environment are greatest.
Continuous and Discontinuous Transition
Two modes have been reported for the prograde transition between saponite and chlorite. The transition has been characterized as a continuous (Bettison and Schiffman,
1988; Bevins, 1991; Chang et al., 1986; Schiffman et al., 1991; Shau and Peacor, 1992;
Son et al., 2001; Weibel, 1999) and discontinuous process (Barrenechea, 2000; Beaufort
et al., 1997; Inoue and Utada, 1991; Inoue, 1984; Jiang and Pecor, 1994; Murakami et al.,
1999; Niu and Yoshimura, 1996; Robinson and Santana de Zamora, 1999; Sandler, 2001;
Schiffman and Staudigel, 1995; Shau et al., 1990). A continuous transformation results
when there is a gradual decrease in the percent of smectite and an increase in randomly
interstratified chlorite-smectite, which eventually attains an ordered 50:50 chlorite-
smectite interstratification. This transition mode is similar to layer stacking sequences 30 observed in smectite-illite subjected to burial diagenesis. On the other hand, discontinuous transformation involves a stepwise decrease in percent smectite independent of chlorite-smectite interstratification (Murakami et al., 1999).
Several studies have suggested that random interstratified chlorite-smectite forms under low grade metamorphism when fluid-rock ratios in the diagenetic setting are low and disequalibrium conditions prevail. For example, Shau and Peacor (1992) used high- resolution transmission electron microscopy (HRTEM) to show that trioctahedral phyllosilicates in hydrothermally altered basalts from DSDP Hole 504B contain poorly defined packets of saponite and mixed-layered chlorite-smectite indicative of crystallization under nonequilibrium conditions resulting from low fluid-rock ratio.
Bettison-Varga and Mackinnon (1997) used HRTEM to show that trioctahedral phyllosilicates in hydrothermally altered basalts from the Point Sal ophiolite contained random interlayerings of smectite- chlorite achieving a 50:50 composition. In this study, randomly interlayered chlorite-smectite represents a metastable phase formed in less- permeable glassy groundmass of the metabasalt where low fluid-rock ratios and less pervasive alteration are implied given the lack of albitization of plagioclase. Chang et al.,
(1986) showed that Cretaceous sedimentary sequences of the northern continental shelf of Brazil contained random interstratified chlorite-smectite to ordered chlorite-smectite
(corrensite) occurred in shales at 70oC and in sandstones at 60oC. The higher degree of
chlorite-smectite ordering is attributed to equilibrium conditions facilitated by increased
fluid flow through the more permeable sandstones.
The discontinuous transition, a dissolution-precipitation reaction, of saponite to
chlorite through corrensite show significant gaps in mineralogy and composition 31 suggesting the transformation process involves discrete steps. Using high-resolution transmission electron microscopy, Murakami et al. (1999) observed corrensite and chlorite growing as discrete domains within saponite without significant amounts of randomly interstratified chlorite-smectite or chlorite-corrensite, suggesting that random interstratification in the conversion reaction is much less stable. Similarly, transmission electron microscopy images taken by Beufort et al. (1997) showed discrete chlorite domains growing within corrensite. In this study, chlorite-corrensite mixed layering was poorly developed in comparison to discrete chlorite and corrensite domains. Inoue and
Utada (1991) studied thermal metamorphism patterns in the volcanoclastic rocks of
Kamikita Japan and found that saponite transformed to chlorite through corrensite with increasing metamorphic grade so that the proportion of smectite layers in the intermediate chlorite-smectite decreased discontinuously in a step-like progression. Chlorite-smectite minerals having intermediate smectite percentages other than those described above were not found. This pattern was also observed in the Pliocene acidic pyroclastic sediments of the Ohyu District in Japan. Inoue et al (1984) observed that expandable layers in chlorite- smectite decreased with depth discontinuously from 100% to 80%, and then from 50% to
40%, and finally 15% to 10% (Figure 2.4). In a subsequent study, Schiffman and
Staudigel (1995) showed that the hydrothermal metamorphism of Pliocene-aged seamount igneous rocks of the Canary Islands La Palma Complex resulted in the transition from smectite to chlorite through a series of depth dependent discontinuous compositional steps between smectite, corrensite and chlorite (Figure 2.5). Furthermore, compositional gaps exist between La Palma smectite, corrensite and chlorite suggesting
32
Figure 2.4: Variation of percentage of expandable layer in chlorite-smectite as a function of depth from drillhole HT-42 of the Ohyu District, Akita Prefecture, Japan (from Inoue et al., 1984).
33
0
-500 Prehnite Chlorite e Pumpellyite Calcit
-1000 Epidote Smectite Garnet Stratigraphic (m) Depth Analcime Corrensite
-1500 Actinolite
Figure 2.5: Mineral zonations of the basement complex La Palma seamount, Canary Islands. Solid line indicates that mineral is nearly ubiquitous and dashed lines indicates that mineral was not present in every sample. Smectite-corrensite and corrensite-chlorite mineral zones overlap suggesting overlap in P-T stabilities (from Schiffman and Staudigel, 1995).
34 precipitation of discreet mineral phases, compositional homogenization and recrystalization with increasing pressure and temperature. Finally, discontinuous transformation of smectite to chlorite through corrensite may be facilitated by high fluid/rock interaction with increasing metamorphic grade (Chang et al., 1986; Schiffman and Staudigel, 1995; Shau and Peacor, 1992).
Chemical Characterization of the Smectite-Corrensite-Chlorite Transition
Analytical data from163 microprobe analyses from ten different studies are used to compare and contrast the chemistry of the smectite-corrensite-chlorite transition between altered mafic igneous and sedimentary rocks. These studies are dominated by phyllosilicate diagenesis during the hydrothermal alteration of mafic igneous rocks
(Schiffman et al., 1991; Schiffman and Staudigel, 1995; Shau and Peacor, 1992) and paleo-hydrothermal environments (Bettison-Varga and MacKinnon, 1997; Bettison et al.,
1991; Bevins, 1991; Inoue and Utada, 1991; Schmidt and Robinson, 1997). While there are few studies involving phyllosilicate diagenesis in hydrothermally altered sedimentary sequences (Barrenechea, 2000; Chang et al., 1986; Hillier, 1993) suggesting this should be an area of future research if the smectite to chlorite transition is to be fully understood.
These systems exhibit a wide range of environmental conditions that can be used to decifer the mechanisms affecting the smectite to chlorite transition.
Smectite and interstratified chlorite-smectite minerals can vary significantly in the number of expandable layers (Drits, 2011; Kogure et al., 2013). For the analysis, samples
containing 0-10% expandable layers are defined as chlorite, samples with 45-55% expandable layers are defined as corrensite and, samples with 90-100% expandable layers are defined as smectite. Chlorite-smectite minerals falling outside the range described 35 above are characterized as interstratified. It is assumed that all phyllosilicates are fully trioctahedral. As such, the formulae presented here were recalculated using the following methodology: 1) balance cation charges on the basis of variable negative charge, 2) assign all of the Si to tetrahedral coordination and add Al to bring the total Si+Al to eight,
3) assign all Al, Fe, Mg, Ti and Mn cations to octahedral sites and, 4) assign the Ca, Na,
K and remaining Mg, after octahedral coordination, to total twelve in the interlayer positions of the swelling component. Since most analytical data represent a continuum between end member smectite and chlorite, formulae were recalculated on a variable oxygen basis in order to compare different trioctahedral phyllosilicates equally. The number of oxygens in the structural formula is determined from an estimate of the proportion of smectite-like and chlorite-like layer types calculated from the sum of octahedral and tetrahedral cations. The following formula, given by Hillier (1993), was used to calculate the percentage of expandable layers in the Chlorite-smectite minerals:
% Expandable Layers = [28(S-20)]/[6(S-28)] where S is the sum of octahedral plus tetrahedral cations in the formula that is cast as chlorite (28 oxygen basis). The assumption of this method is that both smectite and chlorite layers are fully trioctahedral.
The structural formulae for smectite, corrensite, chlorite and interstratified chlorite-smectite from the microprobe analyses show several trends in chemistry. Figure
2.6 is a weight percent ternary diagram of chemical compositions for Si, Al and
Fe+Mg+Mn. Mafic phyllosillicates for both igneous and sedimentary diagenetic environments show a weight percent Si decrease as weight percent of Al and Fe+Mg+Mn
36
Figure 2.6: Ternary plot of %Si, %Al and %Mg+%Fe+%Mn in mafic phyllosilicates demonstrating a decrease in %Si and increase in %Al and %Mg+%Fe+%Mn during the smectite to chlorite through corrensite transformation . All cation values are calculated based on a chlorite-smectite formula with variable number of oxygens. Igneous (Ig) and Sedimentary (Sed) mafic phyllosilicate samples are plotted together for comparison. Sedimentary samples with orange centers represent corrensite (45-55% expandable layers) and those with green centers represent chlorite (0-10% expandable layers).
37 increases in the order smectite > corrensite> chlorite transition. Smectite, corrensite and chlorite plot within distinct regions on the ternary diagram, whereas interstratified chlorite-smectite forms a continuum between smectite and chlorite. Increases in Fe, Mg and Al are attributed to the increase in hydroxide sheet fixation during the prograde transformation of smectite to chlorite.
The conversion of smectite to chlorite through corrensite involves an increase in
Al+3 for Si+4 substitution in tetrahedral sites for these mafic phyllosilicates as well (Figure
2.7). An increasing trend in substitution suggests that the availability of Al+3 may play an important role in whether interstratified chlorite-smectite, smectite, corrensite or chlorite phases predominate (Hillier, 1993). Indeed, Shau and Peacor’s (1992) detailed characterization of hydrothermally altered basalts from DSDP Hole 504B, Leg 83 suggest that well crystallized phyllosilicates typically have relatively low Si/(Si+Al) and are accompanied by significant albitization or zeolitization, which is believed to be the source of Al+3 during phyllosilicate diagenesis. Alteration of basalt without significant
albitization and zeolitization is associated with the occurrence of saponite plus or minus
mixed-layer chlorite-smectite.
The Si content of chlorite-smectite minerals is correlative with the alkali and
alkaline earth content (e.g. Na+, K+ and Ca+2). The Si and Na+K+Ca contents increase in
the order chlorite, corrensite and smectite (Figure 2.8). This trend is linked to the
percentage of smectite-like layers. As the number of smectite-like layers decrease, the
alkali and alkaline earth cation content, which are the principle interlayer cations,
decrease accordingly. This trend suggests that electron microprobe analyses of ‘chlorite’
with relatively high Si+4, Na+, K+ and Ca+2 contents are most likely the result of smectite- 38
3.0
2.5
2.0 [4]Al
1.5
1.0
0.5 0.45 0.55 0.65 0.75 0.85
Si/(Si+Al)
Figure 2.7: Plot of tetrahedral Al total vs Si/(Si+Al) in mafic phyllosilicates demonstrating a strong negative correlation between these two variables. All cation values are calculated based on a chlorite-smectite formula with variable number of oxygens. Symbols are the same as in Figure 2.5.
39
1.20
1.00
0.80
0.60
Na+Ca+K 0.40
0.20
0.00
0.50 0.60 0.70 0.80 0.90
Si/(Si+Al)
Figure 2.8: Plot of interlayer cation totals (Na+Ca+K) vs Si/(Si+Al) in mafic phyllosilicates demonstrating a positive correlation between these two variables. All cation values are calculated based on chlorite-smectite formula with a variable number of oxygens. Symbols are as in Figure 2.5.
40 like interlayering (Bettison-Varga and MacKinnon, 1997; Bettison and Schiffman, 1988;
Shau et al., 1990) . There appears to be no trend in the relationship between Fe/(Fe+Mg) and Si/(Si+Al) (Figure 2.9). A review by Brigatti and Poppi (1984) also suggest that
Mg content increases in the order chlorite › corrensite › smectite, and that Fe/Mg ratios associated with the diagenetic environment regulate the smectite-corrensite-chlorite transformation process when Fe/(Fe+Mg)> 50%, favoring chlorite, but with increasing
Mg content, chlorite transforms into corrensite and then smectite. Schmidt and Robinson
(1997) suggest Fe content increases from smectite to corrensite-chlorite; however, considerable overlap in Fe/(Fe+Mg) exists between corrensite and chlorite. In contrast,
Bettison-Varga and Mackinnon (1997) found no relationship between Fe/(Fe+Mg) and
Si/(Si+Al) in the mafic phyllosilicates from the hydrothermally altered basalt of the
Point Sal ophiolite. Their data show that Fe/(Fe+Mg) is directly controlled by the bulk- rock chemistry. Likewise, Bevins et al. (1991) found variations in Fe/(Fe+Al) in mafic phyllosilicates in metabasites of Wales and Greenland correlate closely with Fe/(Fe+Mg) ratios of the whole-rock (Figure 2.10); however, as shown in Figure 2.11, the same trend does not exist for Si/(Si+Al). Barrenchia et al. (2000) found that chlorite analyses of mudrocks from the Cameros Basin have higher Fe contents than lacustrine carbonates, suggesting a possible link to bulk-rock chemistry and organic content, while overall corrensite had a slightly lower Fe/(Fe+Mg) content than chlorite. Inoue and Utada (1991) and Hillier (1993) documented the transformation of smectite to chlorite through corrensite as being characterized chemically by a constant Fe/(Fe+Mg).
The paucity of analytical data from microprobe studies documenting the chemical characterization of the smectite to chlorite transformation through corrensite makes it 41
0.90
0.80
0.70
0.60
0.50
0.40 Fe/(Fe+Mg) 0.30
0.20
0.10 0.00
0.50 0.55 0.60 0.65 0.70 0.75 0.80 0.85 0.90
Si/(Si+Al)
Figure 2.9: Plot of Fe/(Fe+Mg) vs Si/(Si+Al) in mafic phyllosilicates demonstrating the lack of correlation between these two variables. All cation values are calculated based on chlorite-smectite formula with a variable number of oxygens. Symbols are as in Figure 2.5.
42
0.80
0.70
k 0.60
0.50
0.40 Wales Fe/(Fe+Mg) Bulk Roc Greenland 0.30
0.20 0.30 0.40 0.50 0.60 0.70 0.80
Fe/(Fe+Mg) Phyllosilicate
Figure 2.10: Plot of bulk rock Fe/(Fe+Mg) vs chlorite Fe/(Fe+Mg) for samples from Wales and eastern North Greenland, demonstrating a strong positive correlation between these two variables (from Bevins et al., 1991).
43
0.82
0.80 k 0.78
0.76
0.74
Si/(Si+Al) Bulk Roc
0.72 Wales
0.70 Greenland
0.68 0.50 0.55 0.60 0.65 0.70 0.75 0.80
Si/(Si+Al) Phyllosilicates
Figure 2.11: Plot of bulk rock Si/(Si+Al) vs chlorite Si/(Si+Al) for samples from Wales and eastern North Greenland, demonstrating the lack of correlation between these two variables (from Bevins et al., 1991).
44 difficult for a comprehensive comparison of the similarities and differences between sedimentary and mafic igneous environments. Of the 163 samples analyzed, only 16 were from sedimentary settings. However, even with the limited data there appears to be a similarity in compositional trend: decrease in Na+, K+ and Ca+2, increase in Al+3, Fe+2
and Mg+2 and increase in Al+3 for Si+4 substitution in tetrahedral sites with increasing chloritization.
HRTEM of the Smectite-Corrensite-Chlorite Transition
Insights into the possible mechanisms for interlayer brucitization of smectite in the smectite to chlorite transformation may be inferred from HRTEM studies of biotite chloritization (Veblen and Ferry, 1983) and smectite chloritization(Bettison-Varga and
MacKinnon, 1997; Kogure et al., 2013). Veblen and Ferry (1983) proposed two possible
mechanisms for the development of chlorite from biotite for the hydrothermally altered
granite rocks of south-central Maine: 1) Brucite-like sheet replaces the K-interlayer of
biotite and 2) the loss of tetrahedra from one biotite layer reduces it to a brucite-like
layer. Bettison-Varga and Mackinnon (1997) included a third mechanism to explain the
transformation of smectite to chlorite for the Point Sal ophiolite: Dissolution of 2:1
smectite layer and reprecipitation of two 2:1 units with a brucite-like interlayer in
between. Kogure et al. (2013) results suggest that corrensite precipitated directly from
solution without inheriting smectite structure while the odd number of the successive brucite sheets in the mixed-layer corrensite-chlorite structure along with the similarity of
the polytypic stacking sequences between corrensite and chlorite is evidence of the transition from corrensite to chlorite by the replacement of smectite-like interlayers with
brucite sheets. Schematic representation of the three mechanisms, changes in 45 phyllosilicate layer structure and reaction processes are shown in Figure 2.12. A summary of the processes presented in Figure 2.12 are summarized in Table 2.3. For theformer mechanism of Veblen and Ferry (1983) to be applied to the transformation of smectite to chlorite requires the growth of a brucite-like layer in the interlayer space of smectite resulting in a volume increase. This reaction would require the addition of Mg+2,
+2 +3 + + +2 +4 + Fe , Al and H2O and the loss of Na , K , Ca , Si and H . Modification of the
tetrahedral layers on either side of the brucite-like layer would result in the addition of
Al+3 to tetrahedral sites to accommodate the cell dimension created by the addition of the
brucite-like layer (Bettison-Varga and MacKinnon, 1997; Inoue, 1985; Suquet et al.,
1981). Terminations such as this are found in a variety of phyllosilicates (Olives and
Amouric, 1984; Veblen, 1980; Veblen and Buseck, 1980, 1981), including the smectite-
chlorite series where a brucite-like interlayer appears to terminate into the boundary
between a 10Å and 14Å layer with a resulting increase in volume of the layer structure
(Bettison-Varga and MacKinnon, 1997). For the latter mechanism of Veblen and Ferry to
be applied to the transformation of smectite to chlorite requires the loss of tetrahedra
from either side of the octahedral layer of one smectite layer resulting in a volume
+3 decrease. This reaction requires the addition of Al and H20 in an acidic environment and the loss of Si+4 (substantial amount), Mg+2, Fe+2, Na+, K+ and Ca+2. Veblen and
Ferrry (1983) and Eggleton and Banfield (1985) observed the consumption of mica for
the development of chlorite from biotite; however, this type of termination was not
observed in the alteration of smectite to chlorite by Bettison-Varga and Mackinnon
(1997). The mechanism of Bettison-Varga and Mackinnon (1997).involves the
46
Interlayer water and cations Mechanism 2 Mechanism Chlorite
Smectite 3 Mechanism Brucite-like sheet
Chlorite Chlorite
Brucite-like sheet Brucite-like sheet Mechanism 1Mechanism 1Mechanism 1
Potassium ion
Octahedral sheet
Tetrahedral sheet
Figure 2.12: Schematic representation and process summary of three different chloritization mechanisms for biotite and smectite observed in HRTEM studies: Mechanism 1- the growth of the brucite-like layer is indicated by arrow; Mechanism 2- direction of dissolution of tetrahedral sheets is indicated by arrows; Mechanism 3- direction of dissolution of a smectite layer is indicated by arrow. 46 47
Table 2.3: Summary of the chloritization process for biotite and smectite observed in HRTEM studies.
Mechanism Description Chemical Change Diagenetic Environment HRTEM Study
Mechanism 1 Growth of a brucite-like layer into the Removal of K+ and addition of Hydrothermally altered Veblen and Ferry (1983) (biotite) interlayer region between two mica Mg+2, Fe+2 and Al+3. granite. layers. Volume increase.
Mechanism 1 Growth of a brucite-like layer within Removal of Na+, K+ and Ca+2 1) Hydrothermally altered 1) Bettison-Varga and (smectite) a smectite interlayer. Volume increase. and addition of Mg+2, Fe and Al+3. basalt. Mackinnon (1997), 2) Thermally altered 2) Kogure et al. (2013) sandstone.
Mechanism 2 Formation of a brucite-like layer by Removal of Si+4, Fe+2, Mg+2, Ca+2 Hydrothermally altered Veblen and Ferry (1983) (biotite) removal of the tetrahedral sheets of and K+. granite. one mica layer. Volume decrease.
Mechanism 2 Not observed. (smectite)
Mechanism 3 Not observed. (biotite)
Mechanism 3 Dissolution of a smectite layer and Removal of Na+, K+ and Ca+2 1) Hydrothermally altered 1) Bettison-Varga and (smectite) reprecipitation of a smectite and chlorite and addition of Si+4, Fe+2, Mg+2 basalt. Mackinnon (1997), layers. Significant volume increase. and Al+4. 2)Thermally altered 2) Kogure et al. (2013) sandstone.
47 48 dissolution of a smectite layer and reprecipitation of a chlorite and smectite layer resulting in a substantial volume increase. This reaction requires the addition of Al+3,
Si+4, Fe+2 and Mg+2 and the loss of Na+, K+ and Ca+2. This mechanism is supported by
Meunier et al. (1988), who suggested that during the formation of ordered mixed-layered
clay, the dissolution and redistribution of elements into high-charge saponite layers with
chlorite is necessary to eliminate the crystallographic constraints between the original
smectite and newly formed chlorite. Results from Kogure et al. (2013) based upon
polytypic character suggest that this mechanism best explains the transition of saponite to
corrensite. Bettison-Varga and Mackinnon (1997) observed this most frequently in the
mafic phyllosilicates of the Point Sal ophiolite.
Conclusions
From the data presented in this chapter, the following conclusions may be drawn:
1. Corrensite represents a transitional phase in the Mg-smectite to chlorite transition
and follows the sequence smectite >corrensite >chlorite with increasing pressure
and temperature.
2. The chlorite-saponite form of corrensite is unknown from surface environments
and is, therefore, related to the mesodiagenesis of a precursor mineral or reactions
between Mg-bearing minerals and other phyllosilicates.
3. Recent studies suggest that the transformation of smectite to chlorite through
corrensite is affected by factors such as: 1) temperature; 2) substrate diffusive
capacity and resulting fluid/rock ratios; 3) organic-inorganic reactive interactions
and redox; and 4) whole-rock composition and Al+3 availability. 49
4. Previous studies have shown that the transition may occur in a stepped
progression of smectite to chlorite through corrensite or as a gradual change
involving randomly and regularly interlayered metastable structures associated
with diagenetic microenvironments exhibiting low fluid/rock ratios.
5. Chemically, the conversion of smectite to chlorite via corrensite involves a
decrease in alkali and alkaline earth metals with decreasing Si/(Si+Al), increase in
Al+3 for Si+4 substituition in tetrahedral sites, but variability in the Fe/(Fe+Al) of
octahedral cations which appears to be dominated by whole-rock composition.
6. Three mechanisms have been proposed to explain the transformation of smectite
to chlorite: 1) Precipitaton of a brucite-like layer in the interlayer space of a
smectite resulting in a volume increase in layer structure; 2) loss of tetrahedra on
either side of the 2:1 structure of smectite reducing it to a brucite-like layer
resulting in a volume decrease in layer structure; 3) dissolution of a smectite layer
and subsequent reprecipitation of a chlorite and smectite layer resulting in a
substantial volume increase in layer structure. The former and latter mechanisms
been verified by HRTEM for the chloritization of smectite.
7. Analytical data for the smectite to chlorite transformation in sedimentary systems
are needed. Only with a robust data set combing X-ray diffraction, electron
microprobe and high resolution transmission electron microscopy will a complete
understanding of the mechanisms related to the chloritization of smectite be
illuminated. 50
CHAPTER 3
SEQUENCE STRATIGRAPHIC DISTRIBUTION OF THE CLAY MINERAL
CORRENSITE IN TWO CONTRASTING MISSISSIPPIAN
DEPOSITIONAL ENVIRONMENTS
Introduction
Sequence stratigraphy is a powerful methodology for predicting facies and the depositional control on the spatial distribution of petroleum reservoirs, seals and source rocks
(Posamentier et al., 1988). To date, most sequence stratigraphy studies have utilized lithofacies analysis; however, in recent studies involving siliciclastic sediments, the distribution of diagenetic alteration horizons have been useful in understanding the origin of spatial and temporal variations in the geochemistry of sediments (Al-Ramadan et al., 2005; El-ghali et al.,
2006; Ketzer et al., 2002; Salem et al., 2005; Taylor and Gawthorpe, 2003; Taylor et al., 2004).
This type of geochemical approach has enhanced our understanding of diagenetic processes within a sequence stratigraphic framework, and has created a high-resolution tool for the identification and interpretation of diagenetic processes within stratigraphically significant surfaces and systems tracts (Al-Ramadan et al., 2005; Burns et al., 2005; El-ghali et al., 2006;
Ketzer et al., 2003; Ketzer and Morad, 2006; Ketzer et al., 2002; Salem et al., 2005). This research direction will ultimately help develop more comprehensive and precise models for the distribution of petroleum reservoirs and diagenetic seals (Ketzer et al., 2003). Despite advances in our understanding of these areas, many carbonate studies have focused primarily on the petrophysical aspects of diagenetically altered carbonate (Ehrenberg et al., 2006; Fu et al., 2006; 51
Hopkins, 1999; Petty, 2005; Wierzbicki et al., 2006). Moreover, only a few studies have focused on the diagenesis of carbonate sediments within a sequence stratigraphic framework (Railsback et al., 2003; Saller et al., 1999; Schroeder et al., 2005; Smeesters et al., 2003; Taylor et al., 2000;
Theiling et al., 2007; Westphal et al., 2004). Considering the potential for carbonate rocks to serve as source, seal and reservoir rock for petroleum hydrocarbons, more work is needed to define the significance of diagenetic processes within a high resolution sequence stratigraphic framework over spatial and temporal scales.
Most sequence stratigraphy studies utilizing traditional facies analysis rely heavily upon qualitative methods for the identification of subaerial exposure surfaces. However, sequence boundary identification in carbonates deposited within semiarid-arid settings is often difficult due to the cryptic expression of these surfaces (Railsback et al., 2003; Theiling et al., 2007).
Quantitative methods for the identification of sequence boundaries typically involve the recognition of excursion patterns in δ13C and δ18O values and Sr trace element concentrations in
carbonates collected across stratigraphic horizons. Subaerial exposure surfaces in carbonate
strata have been recognized by relatively low δ13C values below the surface as a result of 12C- rich carbon input from the respiration by and decay of photosynthesizers on overlying land surface (Allan, 1976; Allan et al., 1977; Railsback et al., 2003; Theiling et al., 2007). Changes in
δ18O values have been used to identify subaerial exposure surfaces, where relatively low values
below the exposure are the result of 16O-rich oxygen input from meteoric water and high δ18O values are the result of 18O-rich the evaporative water at the surface. Low Sr concentrations below subaerial exposure surfaces presumably result from the dissolution of aragonite and subsequent removal of Sr2+ in solution. Although stable isotope analysis of carbon and oxygen,
as well as trace element analysis of Sr, have proven to be a useful tool in the identification of 52 subaerial exposure surfaces, replicate sampling along stratigraphic horizons is necessary to accurately identify these surfaces due to intrastratal lateral geochemical variability (Theiling et al., 2007). Replicate sampling at the scale needed to identify subaerial exposure surfaces in carbonates is a labor intensive and costly venture and, therefore, other methods for the identification of these surfaces are needed.
Corrensite as a Geochemical Proxy
Corrensite is a regularly ordered (50:50) interstratified chlorite-smectite clay mineral
(Figure 3.1) shown to be an intermediate phase in the Mg-smectite to Mg-chlorite transition with increasing temperature and pressure. Corrensite is identified in XRD patterns by a strong (001) reflection near 26.5 Å in the air dried state and a shift to 31 Å after glycolation (Figure 3.2). It occurs in two geologic associations: 1) intermediate-mafic igneous rocks altered by hydrothermal fluids; 2) ancient marine evaporites, carbonates or lacustrine rocks. Studies of a modern restricted marine environment suggest the precursor mineral to corrensite, saponite, forms authigenically when pore water Mg/Ca levels rise during the calcite and gypsum precipitation (Hover et al., 1999; Martini et al., 2002). These high ratios are often achieved subsequent to gypsum precipitation in the absence of dissolved inorganic carbon (Hover et al.,
1999), which allows for Mg2+ to be incorporated in the octahedral sites of dioctahedral smectite,
resulting in a Mg-rich trioctahedral smectite. Numerous studies documenting ancient restricted
marine settings have identified corrensite in rock (Andreason, 1992; Bayhan, 2007; Bodine,
1985; Bradley and Weaver, 1956; Droste, 1963; Eardley et al., 1956; Fisher, 1988; Han et al.,
2000; Hillier, 1993; Hints et al., 2006; Jeans et al., 2005; Kopp and Fallis, 1974; Monnier, 1982;
Pay, 2000; Peterson, 1961; Rao, 1992; Ryan and Hillier, 2002; Sandler, 2001; Stein et al., 1991). 55
Figure 3.1: Schematic representation of a corrensite superstructure subjected to different treatments: A) Air dried; B) Ethylene glycol solvation.
53 54
Figure 3.2: XRD patterns of the <2 µm feaction of a sample containing corrensite, illite and minor kaolinite air dried (black) and ethylene glycol sovated (red). Co=corrensite, IL=illite, K=kaolinite. 55
These data suggest that similar processes described by Hover et al. (1990) were occurring in ancient depositional environments. However, since its discovery over 60 years ago (Stephen and
MacEwan, 1951), corrensite has not been considered as a proxy for the identification of sequence boundaries within a sequence stratigraphic framework.
In light of these issues, a study was conducted to test the potential for corrensite as a proxy for sequence boundary identification in carbonate sequences deposited in arid-semiarid marine environments. Two contrasting depositional sequences were chosen for the study: 1) a restricted marine sequence deposited under limited accommodation space; 2) a normal marine sequence deposited under relatively greater accommodation space. Depositional sequence interpretation was based on sequence thickness and traditional facies analysis.
Facies Description
Conventional facies analysis was performed according to in Al-Tawil and Read (2003) and Al-
Tawil et al. (2003), which is consistent with contemporaneous facies analysis described elsewhere (Carney and Smosna, 1989; Ettensohn et al., 1984; Leonard, 1968; Smith et al., 2000;
Smith and Read, 2001; Wynn and Read, 2006). A generalized schematic profile for
Mississippian carbonates of the Appalachian Basin is presented in Figure 3.3. The facies characteristics and their environments of deposition are summarized in Table 3.1.
Photomicrographs of representative facies/rocks reported in Table 3.1 are shown in Figure 3.4.
Sequence Stratigraphy Key Definitions
Sequence stratigraphy is the study of rock relationships within a chronostratigraphic framework of repetitive, genetically related strata bounded by surfaces of erosion or nondeposition, or their correlative conformities (Van Wagoner et al., 1990). A sequence is the fundamental unit in sequence stratigraphy. A sequence is a relatively conformable succession of 56
INNER RAMP MIDRAMP RAMP SLOPE BASIN
0m
20m
KEY
Red Beds Gray shale-siltstone Skeletal GS-PS 60m
Quartz Peloidal Calcareous siltstone Argillaceous skeletal WS-PS GS-PS Laminated Shaly Lime Peloidal GS-PS Sandstone MS-Calcareous siltstone
Pelletal Lime Fine-grained dolomitic Ooid GS-PS WS-MS MS with quartz pseudomorphs after gypsum Paleosol
56 Figure 3.3: A generalized schematic facies profile for Mississippian carbonates of the Appalachian Basin (GS= Grainstone, PS=Packstone, WS=Wackestone and MS=Mudstone) modified from Al-tawil and Read (2003).
57
57
ine sand amp) (Inner ramp) ramp) (Inner crinoids and ostracods
carbonate-clastic ramp facies. carbonate-clastic ramp structureless; and wavy beddedflaser,lenticular locally. containquartz nodules. afterMay pseudomorphs gypsum
Geopetal fill in nodules may be present but rare. fragments. rare. but present be may nodules in pseudomorphs Quartz fill gypsum after fine- grey to dark tan in nodules quartz to white Pink Geopetal to fossiliferous Poorly Quartz peloidal to massive. fissile poorly ordolomite, micrite grained present, If unfossiliferous. Sabkha a pinstriped appearance having in to Light dark gray ramp) (Inner contain Nodules dolomite or may inclusions. anhydrite rounded and Abraded, broken restricted fossils ostracod to eolinite, Coastal mar minor showing dissolution and pitted pitted and slope) fossils ramp marine or dissolution remnant ramp rare slickensides. (Inner with to massive bedded siltstones, ramp) Cross and (Inner sandstone. mollusks, quartz breccia Small showing texture. shaly grading. unfossiliferous. Quartz sandstone grained clay. reverse or to well-sorted to White medium fine green, grey light and displaying Rare quartz silt fine quartz grainstone to include sand ridge, Tidal channel fill Dolomite may and outcrop. abraded peloids some and broken Rounded, skeletal fragments which crystals, dolomite Fine-grained tan. Yellowish abraded ooids, skeletal subangularfinefragments, very and to fossiliferous Poorly ramp) (Inner Flat Tidal mudcracks. arid-semiarid Subaerial, None fibrous and Environment cryptocrystalline Depositional brown, Biota Description to Lithofacies yellow Beds Red White, mudcracks. Caliche siltstones, green and Red,orange, and maroon mudrocks RootBurrow and paleosols,Siliciclastic present. tobe massive Rare may Mottling laminated. green Red, and terrigenous maroon mottled mudrocks paleosols. traces in Root/burrowtraces and to Subaerial marine marginal crusts caliche-coated calcite of Patches fracture and fills. silicified. Variably pisolites. and peloids Subaerial, (Innerhumid r ramp) (Inner and mixed Table 3.1: Carbonate 58
Figure 3.4: Continued.
Figure 3.4: Photomicrographs of facies types for the Mississippian carbonates of the Appalachian Basin: (A) caliche; (B) breccia; (C) siliciclastic paleosol with a highly pitted marine fossil; (D) quartz sandstone; (E) quartz pseudomorph (nodule) after gysum with fortification zoning; (F) quartz-peloidal grainstone (eolinite). 59
Figure 3.4 Continued: (G) microcrystalline dolomite; (H) lagoon pelletal wackestone- packestone; (I) skeletal grainstone; (J) ooid grainstone; (K) midramp argillaceous skeletal wackestone-packstone. 60 genetically related strata bounded by unconformities or their correlative conformities (Mitchum,
1977). Parasequences and parasequence sets are the stratal building blocks of sequences. A parasequence is a relatively conformable succession of geneticallyrelated beds or bedsets bounded by flooding surfaces or their correlative surfaces (Van Wagoner et al., 1990; Van
Wagoner et al., 1988). Parasequences form distinct stacking patterns within a sequence and are referred to as parasequence sets. Parasequence sets can display aggradational, retrogradational or progradational stacking patterns (Figure 3.5).
Sequences are defined by the lower bounding surface. The two types of bounding surfaces for sequences are referred to as type-1 or type-2. A type-1 sequence boundary (SB) forms when the rate of eustatic sea-level fall exceeds the rate of subsidence at the depositional- shoreline break, producing a relative fall in sea level at that position (Posamentier et al., 1988;
Van Wagoner et al., 1990; Van Wagoner et al., 1988). During the development of a type-1 SB, sediments become subaerially exposed, forming an unconformity landward of the depositional- shoreline break. A type-2 SB is interpreted to form when the extent of the sequence-bounding
Unconformity reaches seaward to the position of the previous shoreline. During the development of a type-2 SB, the sediments do not experience exposure seaward of the high stand systems tract.
A sequence can be subdivided into systems tracts, which are defined by their position within the sequence and by the stacking patterns of parasequence sets and parasequences (Van
Wagoner et al., 1988). The term systems tract is used to designate three subdivisions within each sequence. A type 1 sequence contains lowstand, transgressive and high stand systems tracts and type 2 sequence contains shelf margin, transgressive and highstand systems tracts (Figure 3.6).
The lowstand systems tract (LST) is bounded below by the SB and above by the first major 61
Figure 3.5: Cross-section of parasequence-stacking patterns in parasequence sets (modified from Van Wagoner et al., 1990). Four shallowing upward parasequences are shown in each cross sectional view. 62
Type-1 Sequence
Type-2 Sequence
Figure 3.6: Stratal patterns and systems tracts in sequences: Type-1 sequence and; Type-2 sequence (from Van Wagoner et al., 1988). 63 flooding surface called the transgressive surface (Posamentier et al., 1988; Van Wagoner et al.,
1988). Typically the LST is missing in carbonate sequences due to the dissolution of carbonateand the lack of siliciclastic input characteristic of these settings. Above this first major flooding surface marks the beginning of the transgressive systems tract. The transgressive systems tract (TST) contains a retrogradational parasequence set that progressively deepens upward and is bounded above by the maximum flooding surface (MFS). The MFS is represented by the deepest water facies and represents the greatest landward extent of shoreline transgression. The highstand systems tract (HST) is bounded below by the MFS and above by the next SB, which is the beginning of the next sequence. The early HST commonly consists of an aggradational parasequence set; the late HST is composed of one or more progradational parasequence sets (Posamentier et al., 1988; Van Wagoner et al., 1988). In type-2 sequences, shelf-margin systems tract (SMST) is defined by weakly progradational-aggradational parasequence set at the end of the HST.
For sediments to accumulate in a basin there must be space available below some level
(base level) above which erosion will occur. This space made available for potential sediment accumulation is referred to as accommodation space and is a function of both eustatic sea level and subsidence. Subsidence can occur by several processes including overthrust loading and viscoelastic relaxation, lithospheric stretching at divergent margins, lithospheric cooling or sediment compaction. For example, if eustatic sea level is falling at a certain rate and sea floor subsidence matches that rate, then relative sea level remains unchanged and no new accommodation space is made available for sediments to deposit. However, if eustatic sea level 64 is falling more slowly than sea floor subsidence, the net effect will be a relative sea-level rise and new space will be added. Therefore, changes in relative sea level reflect the combined effect of sea-level change and subsidence.
Sequence Thickness and Accomodation
Sequence thickness is a function of the rate of accommodation generated by tectonic and eustatic mechanisms as well as sedimentation rate. Since accomodation creates space for sediments to deposit and the rates of carbonate deposition typically equal or exceed the rate of accommodation, the thickness of a particular sequence is a function of the available accommodation space during a particular period of deposition. During the Mississippian, low subsidence rates of the distal foreland of the Appalachian Basin resulted in a highly thinned sequence succession (Al-Tawil and Read, 2003). In contrast, proximal foreland deposits of the
Mississippian show a pattern of sequence thickening which suggests an increase in amplitude of sea-level change during Upper Meramecian to Chesterian (Al-Tawil and Read, 2003; Al-Tawil et al., 2003). The increase in sequence thickness is attributed to the onset of global icehouse conditions, causing the ramp to be flooded to depths of only 10 meter initially to tens of meters in the later Chesterian (Al-Tawil and Read, 2003).
Geologic Setting of North America during the Mississippian Period
During the late Mississippian period, the northern edge of eastern Tennessee was situated at approximately 28 degrees south latitude (de Witt and McGrew, 1979; Scotese and Mc Kerrow,
1990) within the dry trade-wind belt (Figure 3.7). By the early Pennsylvanian, the northward movement of the North American plate placed the area in the wet equatorial belt (de Witt and
McGrew, 1979; Scotese and Mc Kerrow, 1990). This gradual shift from a semi-arid condition during the Meramecian and earliest Chesterian epochs to a wetter climate by late Early 65
A)
B)
Figure 3.7: Paleogeographic reconstruction; A) Late Mississippian time showing that the position of Georgia and Tennessee within the dry trade-wind belt and; B) early Pennsylvanian time illustrating the movement of the Appalachian Basin into the wet equatorial belt (Scotese, 2000). 66
Chesterian time is recorded in pedogenic features associated with subaerial exposure surfaces
(Ettensohn et al., 1988). The early arid-semiarid climate was a major factor responsible for caliche formation within paleosols; however, by the Early Chesterian time, the caliches were partially destroyed by dissolution, creating a leached, clayey residual soil on top of earlier caliche soils (Ettensohn et al., 1988).
The study areas for this project are located within the Mississippina Appalachian foreland basin. During this time, the Appalachian basin consisted of an eastern arcuately elongate miogeosynclinal segment extending from eastern Pennsylvania to northern Mississippi (de Witt and McGrew, 1979). Late Paleozoic tectonism in the Appalachian orogen is interpreted to be the result of a linear contractional orogenic belt which formed during the collision of North America and African-South American continents (Scotese and Mc Kerrow, 1990). The collision culminated in a series of large scale cratonward thrusting of Paleozoic rocks onto the North
American craton (Mack et al., 1983). Thrust-sheet loading onto the craton produced an elongate trough cratonward of the orogen and a westward migrating peripheral bulge on the distal margin
of the basin during the loading phase of the Acadian tectophase (Figure 3.8). Later, a regional Middle Mississippian (Osage-Meramecian) unconformity and associated shallow- marine sediments suggest an eastward bulge movement accompanying load relaxation
(Ettensohn, 1993). However, the many-fold thickening of strata eastward in the basin and the presence of upper Meramecian to Chesterian deep water facies in the proximal foreland of
Kentucky suggest that regional subsidence patterns cannot be explained exclusively by the basinward migration of an anti-peripheral bulge during unloading (Al-Tawil and Read, 2003).
67
Figure 3.8: Foreland deformation model in which basin subsidence, peripheral upwarping and forebulge migration (white arrows) are a response to thrust-belt loading (black arrow): 1) Lithosphere responds elastically to thrust-loading and forebulge migrates away from load; 2) lithosphere adjusts to loading with basin deepening and accentuated upwarping and change in forebulge migration direction; 3) viscoelastic relaxation resulting in progressive uplift of forebulge and its migration towards the load. Triangles mark the position of forebulge at different times of foreland development (After Quinlan and Beumont, 1984).
68
Consequently, Al-Tawil. and Read (2003) suggest reinitiation of compression and subsequent thrust loading, possibly the result of the Quachita orogeny to the southwest, as a mechanism for rapid subsidence in the Tennessee-Kentucky foredeep.
The Cincinnati Arch, interpreted to be the peripheral bulge of the Appalachian foreland basin, formed initially during the Taconic tectophase (Ettensohn, 1992; Pope et al., 2009)The
Cincinnati and Waverly Arch were topographic highs that were active during the Meramecian and Chesterian time (Dever, 1999; Ettensohn, 1980; Ettensoln, 1975; Pryor and Sable, 1974;
Sable and Dever, 1990; Woodward, 1961). The Jessamine Dome in central Kentucky and the
Nashville Dome in central Tennessee, situated along the long axis of the Cincinnati Arch, are separated by a passageway between the Appalachian and Illinois Basins referred to as the
Cumberland Saddle. Faults, arches and domes active during the Mississippian resulted in rapid thickness changes across down dropped blocks with thinning and erosion over highs (Al-Tawil,
1997).
Regional Stratigraphic Framework
The stratigraphy and biostratigraphy of the Mississippian system in Georgia are given in
Hayes (1891, 1894, 1902), Butts and Gildersleeve (1948), Allen and Lester (1954), Cressler
(1970), McLemore (1971), Dutro et al. (1979), Thomas and Cramer (1979), Rich (1982a, b,
1986), and Rich and Algeo (1982). In northwest Georgia, the carbonate and siliciclastic rocks of the Mississippian System range in thickness from approximately 360 to 460 m (Thomas and
Cramer, 1979). These rocks are mainly carbonates and they are divided into five formations: 1) the Fort Payne and 2) Tuscumbia Formations (both Meramecian); 3) and the Monteagle, 4)
Bangor and, 5) Pennington Formations (all Chesterian). The bedded chert of the Fort Payne 69
Formation is easily distinguishable from the nodular chert of the Tuscumbia Formation. The contact between the Tuscumbia and overlying Monteagle Formations is arbitrarily placed between a bioclastic limestone-nodular chert unit of the Tuscumbia Formation and the noncherty bioclastic-oolitic limestone of the Monteagle. The Hartselle interval, a siliciclastic unit, divides the Monteagle and Bangor Formations. The Bangor Formation is composed of gray and dark- olive ooid grainstone, skeletal wackestone-packstone and spiculiferous wackestone-packstone.
The Bangor Formation grades upward into the gray calcareous shale and maroon-green mudstone of the Pennington Formation.
These carbonate facies pass gradationally into a clastic facies that is over 750 m thick towards the southeast. Details of the transition between the two facies are lacking owing to the removal of Mississippian rocks by erosion along the Peavine anticlinorium (McLemore, 1971;
Thomas and Cramer, 1979). However, limestone tongues within the southeastern clastic units suggest a lateral transition distinguished by intertonguing of distinct facies (Rich, 1986; Thomas and Cramer, 1979). The evident facies pattern across northwest Georgia suggests that a delta system carried siliciclastic sediments from the eastern Appalachian highlands north and westward to intertongue with a northwest carbonate facies of the Appalachian Basin, which was deposited on a shallow ramp.
The rocks of the study interval are exposed in an abandoned quarry on the southeast edge of Lookout Mountain in Walker County, Georgia (Figure 3.9). The Tuscumbia and Monteagle
Formations make up the study interval and consist of lagoonal micritic carbonate, ooid and skeletal grainestone-packstone shoal complexes and bryozoan-rich subtidal packstone
70 wackestone. Although the sequence stratigraphy for the Mississippian system in Georgia has not been defined on a regional scale, sequence stratigraphic nomenclature has been preliminarily assigned (Bulger, 2010) to the unconformity bound series of rocks exposed in the quarry for purposes of comparison.
The stratigraphy and biostratigraphy of the Mississippian system for northeastern
Tennessee are given in Campbell (1893), Butts et al. (1925), Englund and Smith (1960),
Englund (1964, 1968), Neuman and Nelson (1965), Mixon and Harris (1971), Hasson (1972),
Milici (1974), Pryor and Sable (1974), Craig and Varnes (1979), de Witt and McGrew (1979)
Dutro et al. (1979), Milici et al. (1979), Maples and Waters (1987), Dever (1995) and Al-Tawil and Read (2003). The Mississippian system is composed largely of carbonate rocks that were deposited on a relatively stable platform to the west, and of terrigenous clastic and carbonate rocks that were deposited in a subsiding geosyncline to the east (Milici et al., 1979). On the northern tip of the Cumberland Plateau in Tennessee, the system is over 1000 ft. thick and thickens toward the east to over 6500 ft. near the Clinch Mountain Belt.
The lower most Mississippian unit, the Grainger Formation, consists mostly of carbonaceous and pyritic grayish-black shale and greenish-gray, greenish-red shale with abundant siderite nodules. The Grainger Formation reaches a maximum thickness of 320 m near
Chilhowee Mountain and progressively thins eastward to approximately 70 m on the northwest side of the Pine Mountain block (Millici et al., 1979). It grades laterally into the Ft. Payne
Formation, which extends from the western part of the Valley and Ridge, beneath the
Cumberland Plateau to the Highland Rim (Millici, 1979). The Fort Payne Formation is 71 composed of several lithologies in Tennessee which include: finely crystalline bedded cherty dolomite, a heterogeneous mixture of carbonate and terriginous clastic material, a calcareous shale and siltstone, and a cherty argillaceous limestone.
The Newman Limestone consists of those beds between the top of the Ft. Payne-Grainger
Formation and the base of the Pennington Formation. East of Clinch Mountain, it is estimated that the Newman Limestone reaches a thickness between 637 and 914 m (Milici et al., 1979).
The group thins toward the west attaining a thickness of approximately 223 m along Cumberland
Mountain to no more than 210 m in Elk Valley (Milici et al., 1979). The Newman Limestone is subdivided into ten formations for the eastern Pine Mountain outcrop belt and is described in ascending order as follows: the St. Louis, Saint Genevieve, Warix Run/Paoli, Beaver Bend,
Reelsville, Beach Creek Cave Branch, Haney, Hardinsburg, Glen Dean and Reynolds Formation
(Al-Tawil and Read, 2003; McFarlin and Walker, 1956). The Chesterian Saint Genevieve, Warix
Run/Paoli, Beaver Bend and Beach Creek, all Chesterian, were deposited during arid-semiarid conditions in which a moderate fourth-order sea level rise favored oolite deposition. However, by mid to late Chesterian, the onset of a more humid climate coupled with an increased magnitude of fourth-order sea level change resulted in the deposition of open marine skeletal limestones bounded by marine siliciclastic units of the Haney, Hardinsburg, Glen Dean and
Reynolds units.
The Newman Group is succeeded by the Pennington Formation which is composed of dolomite, limestone, red, green or gray shale, fine-grained sandstone and conglomeratic sandstone (Campbell, 1893). In general, the Pennington Formation is thicker and contains a greater proportion of terriginous clastic deposits on the eastern side of the Cumberland Plateau, while to the west it is thinner and more calcareous (Milici et al., 1979). Deposits of the 72
Figure 3.9: Thickness of Mississippian sediments in the Appalachian Basin (modified from de Witt, 1975). The triangles identify Lookout Mountain (green) and Pine Mountain (blue) outcrop localities.
73
Pennington Formation in Tennessee have been interpreted as littoral-nondeltaic, and include tidal flat, tidal channel, levee and intertidal deposits (Bergenback et al., 1972), and offshore sandbars (Milici, 1974). Overall, the Pennington ranges in thickness from 30 to 152 m (Milici et al., 1979).
The rocks of the study interval are exposed along the north end of a road cut along Highway 75 near Jellico, Tennessee (Figure 3.9). The study interval is within the St. Louis Formation and makes up the M1 sequence defined by Al-tawil and Read (2003). Rocks of the M1 sequence are dominated by restricted inner ramp facies such as yellowish tan dolomitic tidal flat complex and greenish quartz sandstone, representing a restricted clastic shoreline complex, containing quartz pseudomorphs after gypsum. However, a skeletal-peloidal packestone unit representing a normal marine setting does exist in the middle of this sequence.
Methods
Fifty hand samples were collected from a 12 meter vertical transect from an abandoned quarry on the southeast flank of Lookout Mountain 16 miles northwest of LaFayette on Highway
136 in Walker County, Georgia. Eighteen samples were collected from a 5 meter section of a sequence exposed along the Pine Mountain overthrust 2.5 miles southeast of Jellico on
Tennessee Highway 75 in Campbell County, Tennessee. Fist-sized samples were collected with a rock hammer at the base, middle and top of parasequences and where facies transitions occurred.
Thin sections were prepared from each sample to classify and identify facies relationships according to Dunham (1962).
Each sample was crushed and approximately 20 g of <2 mm fraction was isolated for clay mineral analysis. Clay minerals were separated from carbonate rock with a buffered acetic 74 acid solution (pH 4.5), repeatedly washed with deionized water, and then disaggregated in a solution of deionized water and hexametaphosphate. The <2 µm fraction was separated from the slurry by centrifugation. Clays were concentrated by evaporating the liquid obtained during centrifugation. The dried clay fraction was resaturated with approximately 5 ml of water, disaggregated and sonified for approximately 10 seconds. An oriented clay mineral slide was prepared by pipetting the clay slurry onto a petrographic slide and allowing the slide to dry.
Oriented clay slides were analyzed for clay mineralogy with a Bruker D8 Advance XRD using the following parameters: 2 to 35o 2 Θ, 40 ma, 40 kv at 5o per minute step and 0.02 increment.
Clay mineral slides were run both dry and ethylene glycol saturated.
The presence of quartz, calcite and dolomite was quantitatively assessed with XRD. A
bulk sample was lightly crushed. Approximately 8 g of the less than 2 mm fraction was placed in
a McCrone Mill with ethanol and ground for 10 minutes. Randomly oriented bulk powders were
analyzed with a Bruker D8 Advance XRD under the following parameters: 2 to 75o 2 Θ, 40 ma,
40 kv at 5o per min step and 0.02 increment. Percentages for quartz, calcite and dolomite were calculated with a Rietveld refinement algorithm provided in the program Topas ® by Bruker.
Results
Conventional facies analysis of the Lookout Mountain section revealed a series of higher order shallowing upward cycles (parasequences) superimposed upon a larger scale deepening and then shallowing trend bounded by subaerial exposure surfaces (sequence). The lower bounding surface of the sequence has numerous sedimentary structures suggesting subaerial exposure: truncation, dessication cracks, dense micrite, restricted fauna dominated by ostracods
and chert pseudomorphs after anhydrite (Figure 3.10a), while the upper bounding surface was
more cryptic and represented by an undulatory contact succeeded by an unfossiliferous 10 cm 75
Figure 3.10: Sequence boundaries of the Lookout Mountain sequence: A) The lower bounding surface with numerous quartz pseudomorphs after anhydrite-gypsum nodules, truncation features and mudcracks and; B) the upper bounding surface consisting of a 10 centimeter oxidized horizon which is much more cryptic in physical expression. 76 oxidized carbonate shale horizon (Figure 3.10b). Based upon facies analysis, fossil content and sequence thickness, sufficient accomodation space was created to produce a sequence where open marine conditions prevailed. Figure 3.11 is a key to interpret the symbols presented in stratigraphic sections. A total of eight parasequences were recognized in the 10 meter thick sequence exposed in the quarry at Lookout Mountain (Figure 3.12). Within the transgressive systems tract (TST) of the sequence, the carbonate rocks transition from a restricted inner ramp facies (parasequence P1) to a midramp shoal (parasequence P2) complex and, finally an upper ramp slope facies (parasequences P3-P5). There are four parasequences identified in the TST.
With the exception of P1, parasequence stacking patterns exhibit a coarsening upward pattern (P2-P4), which progress toward deeper water facies with each succeeding parasequence. Parasequences are bounded by planar flooding surfaces succeeded by fine-grained wackestone, muddy packstone or packstone which grade conformably into coarser grained skeletal and ooid grainstones. Parasequences P3-P5 represented by an aggrading bryozoan-rich wackestone-packstone parasequence set. The flooding surface of P5 is interpreted as the maximum flooding surface (MFS) for the sequence on the basis that it is represented by a quartz concentration maximum within the ramp slope facies. Stacking of articulated fronds of fenestrate bryozoans, semi-articulated columnals and calyxes of crinoids and brachiopods suggest a low energy, normal marine upper ramp slope environment (Feldman et al., 1993); however, the presence of pyrite and lack of bioturbation suggests anoxic conditions most likely prevailed below the water-sediment interface.
The overlying high stand systems tract (HST) is represented by the transition of the upper ramp slope facies (parasequences P5 and P6) to midramp oolitic shoal facies (parasequences P7 and
P8). There are four parasequences in the HST. Once again, parasequences exhibit a shallowing 77
Figure 3.11: Interpretive key for the stratigraphic sections presented in Figure 3.12 and Figure 3.13.
78
Figure 3.12: Sequence stratigraphy and mineralogy of the Look Out Mountain locality in Georgia. Corrensite is present near the position of subaerial exposure surfaces where restricted marine conditions would have prevailed in a semi-arid/arid climate. Subsequent alteration of pre- existing clay minerals by Mg-rich hypersaline reflux would result in the neoforamation of corrensite precursor minerals. Burial diagenesis would alter precursor minerals to corrensite. Mixed-layered illite-smectite is present in open marine facies. 79 upward trend. Flooding surfaces are overlain by wackestone, which coarsens upward into fine- grained skeletal grainstone and finally an oolitic grainstone. The top of parasequence P8 of the
TST and early HST deposits contained smectite, illite and minor kaolinite. Within the shallowing phase of the HST, corrensite appears progressively deeper within each succeeding parasequence cycle (parasequences P7 and P8). This trend suggests marine water became increasingly restricted during the last two phases of parasequence development within the HST is the upper bounding unconformity for the sequence prior to subaerial exposure (Figure 3.12).
Percent quartz-dolomite-calcite data show trends that generally follow parasequence and sequence development. The bases of parasequences have lower percent calcite values which progressively increase as the percentage of carbonate allochems increases within each shallowing upward parasequence cycle. The increase in quartz at the base of parasequences is most likely a function of decreased carbonate production by benthic organisms, which is easily disturbed by increased water depth (Sarg, 1988; Schlager, 1981). During the overall development of the sequence, the percent calcite decreased during the development of the TST
(deepening phase), attained minimum value at the maximum flooding surface (MFS) and increased with the development of the HST (shallowing phase). In addition, parasequence thickness (P1-P4) decreases towards the MFS. These trends suggest that accommodation space outpaced carbonate production and sedimentation during the development of the TST. The decrease in percent calcite and concomitant increase in dolomite and quartz towards the top of the sequence (the top of parasequence P8) corresponds with a subaerial exposure and an oxidized paleosol representing the upper sequence boundary for the sequence.
Facies analysis of the Pine Mountain section identified three parasequences in the M1 sequence of the Pine Mountain section (Figure 3.13). Parasequences exhibit fining upward cycles 80
80 te-
of this sequence resulted in Tennessee. Corrensite and poorly ordered ls would have been optimal. Mixed-layered illi ls would have been optimal. omodation space during the deposition omodation logy of the Pine Mountain locality in e the neformation of corrensite precursor minera e the neformation corrensite appear throughout the entire sequence. Limited acc Limited sequence. the entire corrensite appear throughout restricted marine conditions wher above. the sequence tract of systems does not occur until the transgressive smectite Figure 3.13: Sequence stratigraphy and minera 81 that were deposited under minimal accomodation space and dominated by restricted inner ramp facies. Sediments were either cemented with or composed of fine-grained dolomite. The majority of beds in this sequence lacked fossils with the exception of a packstone bed in the middle of the sequence, which was composed of poorly sorted, peloids, echinoderm, bryozoan, brachiopod fragments and sub-angular quartz grains. The faunal diversity and fossils in this unit is indicated waters of normal marine salinity in comparison to the bounding unfossiliferous dolomitized mudstones. As such, this section represents a lagoonal facies based upon thin section analysis and was identified as the MFS within the sequence. In thin section, quartz nodules in sandstone at the base of the sequence show the presence of length slow quartz, fortification zoning and anhydrite inclusions, which are evidence for quartz replacement of anhydrite-gypsum nodules
(Chowns and Elkins, 1974; Elorza and Rodriguez-Lazaro, 1984; Folk and Pittman, 1971;
Gomez-Alday et al., 2002; Maliva, 1987; Milliken, 1979; Ulmer-Scholle et al., 1993).
Considered collectively, this evidence suggests the sandstone was deposited in a restricted
Clay mineral analysis of the Pine Mountain samples revealed the presence of corrensite and poorly ordered corrensite throughout entire sequence and the absence of corrensite within the deepening phase of a second sequence. The skeletal-peloidal packstone unit that separates the
TST and HST contained poorly ordered corrensite and illite. This interval is interpreted to be the maximum flooding zone based on facies analysis. Corrensite is absent directly above the sequence boundary-transgressive surface of the second sequence (C1 of Al-Tawil and Read,
2003), and is replaced by poorly ordered smectite-chlorite, chlorite, illite and minor kaolinite.
Poorly ordered corrensite in one sample immediately above the transgressive surface of the C1 sequence most likely represents a transgressive lag deposit where reworked pre-exposure sediment was mixed with post-exposure sediment.
82
Trends in percent quartz-dolomite-calcite data for the Pine Mountain locality were poorly developed within each parasequence. Overall, calcite production from biogenic sources was limited throughout most of the sequence with the exception of the skeletal-peloidal packstone unit. Quartz-rich sandstones represent an upper shoreface sand ridge that fines upward into a more dolomite-rich tidal flat facies. The first two parasquences show a pattern in % carbonate that is similar to the Lookout Mountain parasquences. The base of the third parasequence is a skeletal-peloid packstone, which is dominated by calcite grains. The base of this parasequence contained a high percentage of calcite that dropped as the facies transitioned from skeletal- peloidal packstone to dolomite-siliciclastic-rich tidal flat facies.
Discussion
Parasequence development is markedly different between the Lookout Mountain and
Pine Mountain localities and reflects of contrasting modes of deposition and accommodation.
The sequence at Lookout Mountain is close to 12 m thick, which is approximately 3 times thicker than the M1 sequence at the Pine Mountain locality. Sufficient accommodation space was available to support coarsening upward cycles dominated by normal salinity marine fauna.
Parasequence stacking in the TST follow a predictable deepening upward pattern starting with inner ramp lagoon, transitioning to mid ramp shoal and finally upper ramp slope facies. The parasequences of the HST shallow upward, with the lowermost unit being a ramp slope facies followed by a shoal facies and capped by a subaerial exposure surface. The poorly fossiliferous shaly grey limestone at the base of the parasequences is the result of interruption in carbonate production during the deepening phase of each parasequence. In contrast, the analogous M1 sequence at Pine Mountain is composed of restricted marine inner slope facies highlighting the limited accommodation created during deposition. All of the parasequences exhibited a fining
83 upward trend and were capped by a tidal flat facies. The sandstone in the lower two parasequences are shallow marine facies that shallow upward into a tidal flat facies. The pelletal packstone unit in the third parasequnce is characteristic of a lagoonal facies and represents a maximum water depth.
Data from the Pine Mountain and Lookout Mountain localities suggest a link between clay and whole rock mineralogy, and sequence development. Corrensite and dolomite are ubiquitous components of the whole rock mineralogy in restricted marine facies associated with sequence boundaries. This is not surprising considering micritic dolomite is common in sediments of modern intertidal-supratidal sabkhas (Chafetz and Rush, 1994; Illing and Taylor,
1993; Illing and Wells, 1964; Mackenzie, 1981; Shinn, 1983; Warren, 2000; Warren, 1991;
Wenk et al., 1993) as well as Mg-rich precursors to corrensite (Hover et al., 1999; Martini et al.,
2002). A TEM study showed that much of the dolomite in the Abu Dhabi sabkha is precipitated directly from solution rather than being a replacement of precursor aragonite (Wenk et al., 1993)
. The presence or absence of dolomite is most likely related to factors that constrain the formation of dolomite itself. Baker and Kastner (1981) demonstrated experimentally that the rate of dolomitization is increased when the level of dissolved sulfate is decreased. They concluded that high levels of sulfate, such as in normal sea water, inhibit dolomite formation. In contrast, low sulfate levels in marine pore fluids that have undergone microbial reduction may create suitable environments for dolomite precipitation (Warren, 2000). Dolomite precipitation is also enhanced by the presence of dissolved organic carbon (DOC) in the sediment (Bultler, 1969) while diminished in DOC poor environments (Hover et al., 1999).
Normal marine facies have a mineralogical composition that is easily distinguishable from restricted marine facies in the two study sites. As the former is characterized by calcite,
84 illite-smectite, illite and kaolinite. Calcite is probably derived from the skeletal remains of marine organisms and pore-filling diagenetic calcite while the clay mineral suite is most likely derived from the thermal alteration of detrital smectite, illite and kaolinite washed into the basin from the Appalachian highlands. The change in mineralogy to restricted marine facies is abrupt and easily recognizable with the presence of dolomite at the transgressive surface between sequences and the presence of dolomite and corrensite immediately above the sequence boundary.
Corrensite as a Proxy for Sequence Boundary Position
The results suggest corrensite may be a helpful in recognizing the position of sequence boundaries in carbonate rocks deposited in basins during arid-semiarid conditions. At the
Lookout Mountain locality, corrensite appears in relatively narrow (meter scale and less) bands in close association with restricted marine inner ramp facies below sequence boundaries. The occurrence of corrensite at the top of shoaling upward cycles of the HST suggest hypersaline conditions may have existed with the development of intershoal lagoons. Reflux of hypersaline water through shoal sediments could have facilitated clay mineral diagenesis towards Mg-rich corrensite precursor minerals. The relatively narrow intervals in which corrensite appears suggest that hypersaline reflux was limited to the uppermost layers of sediment during limited accommodation space.
The distribution of corrensite in the restricted marine sequence at the Pine Mountain locality is further evidence for the utility of corrensite to identify sequence boundaries in rocks deposited during arid-semiarid conditions. Corrensite appears predictably in restricted marine
85 facies which dominated this sequence. The absence of corrensite in the skeletal-peloidal packstone, which is the maximum flooding event for this sequence, suggests the appearance of corrensite is constrained to restricted marine facies that experience hypersaline reflux.
Confirmation of Facies Interpretation
Sequence stratigraphic analysis utilizing outcrop and core data depends heavily upon accurate identification and interpretation of facies. However, the qualitative nature of facies analysis is at times subject to misinterpretation. The presence of corrensite in carbonate rock could be used to constrain facies interpretation. In the Mississippian for example, muddy carbonates have been interpreted as low energy subtidal lagoon facies (Choquette and Steinen,
1980), as fore-shoal facies (Dodd et al., 1996), as facies formed both landward and seaward of shoals (Cluff and Lineback, 1981), and in low-energy environments landward of shoals, in swales and channels between and in front of shoals (Smith et al., 2000). Likewise, similarities in physical expression between tidal flat and deeper-water facies could lead to erroneous interpretations. Cores with thin beds of massive dolomitized pelletal packstones, formed in low- energy subtidal lagoons lacking peritidal features such as mudcracks or microbial laminations, are difficult to distinguish from deeper-water calcisiltites (Khetani and Read, 2002). Results from this study suggest that the two facies could be easily distinguished geochemically. The ramifications resulting from the misinterpretation only highlights the need for quantifiable methods of facies analysis.
Conclusions
This preliminary study highlights the utility of the clay mineral corrensite to serve as a proxy for sequence boundary identification in carbonate rocks deposited during arid-semiarid
86 conditions. Furthermore, the clay mineral corrensite may serve useful as a climate proxy if analyzed along temporal scales. From the data presented in this study, the following conclusions can be drawn:
1. Surfaces of subaerial exposure in carbonates formed during periods of high aridity
have unique mineralogy that is recognizable with X-ray diffraction methods.
2. Corrensite appears below sequence boundaries in carbonates deposited during arid-
semiarid conditions. Corrensite does not appear to dominate in rock where normal
marine conditions prevail. The ubiquitous presence of corrensite below a suspected
boundary and subsequent disappearance above is quantifiable geochemical evidence
for a sequence boundary.
3. Linking sediment diagenesis to a sequence stratigraphic framework can provide
valuable information concerning temporal distribution of diagenetic alterations in
clay minerals.
4. Corrensite may be used to constrain facies interpretations in carbonate sequence
stratigraphic studies. The presence of corrensite in carbonate rock is strong evidence
87
CHAPTER 4
GEOCHEMICAL AND MINERALOGIC EXPRESSION OF SUBAERIAL EXPOSURE
IN MISSISSIPPIAN CARBONATES OF THE APPALACHIAN BASIN: EVIDENCE
FROM THE PINE MOUNTAIN OVERTHRUST BELT IN NORTH CENTRAL
TENNESSEE
Introduction
Sequence stratigraphy is a powerful methodology for predicting the depositional control on the spatial distribution of facies. The hierarchial nature of sediment packages characterized by this approach helps identify chronostratigraphically significant surfaces which may be used to define stratigraphic sequences (Posamentier and Vail, 1988; Vail and Mitchum, 1977; Van
Wagoner, 1995; Van Wagoner et al., 1988). Sequence stratigraphy provides guidelines for the interpretation of depositional systems under varying tectonic and eustatic regimes within a chronostrigraphic context. In regards to carbonates, the shallow-water nature of most of these depositional systems makes them susceptible to rapid sea-level change which can potentially create omissions in stratal hierarchies and systems tracts (Al-Tawil and Read, 2003; Al-Tawil et al., 2003; Smith and Read, 1999). In addition, subaerial exposure surfaces can often be cryptic in their physical expression due to the lack of terrestrial organisms that facilitate pedogenesis
(Saller et al., 1994; Smith and Atkinson, 1976). Since rates of diagenetic alteration of metastable sediment increase greatly with increased rainfall and recharge, exposure surfaces may be difficult to recognize during periods of arid-semiarid climate when net moisture deficit result in low fluid flux per unit time (James and Choquette, 1984; Smart and Whitaker, 1991; Ward,
88
1973). Therefore, additional markers that can be used to identify subaerial exposure surfaces can further enhance our understanding of the processes generating sequence architecture.
The interpretation of carbonate systems within a sequence stragraphic framework can be aided by geochemical techniques (Railsback, 1993; Railsback et al., 2003; Saller et al., 1999;
Taylor et al., 2000; Theiling et al., 2007). Carbonate systems are unique depositional settings where controls on sedimentation are essentially biological (Schlager, 1981). Carbonate producing organisms keep pace with most subsidence controls (Al-Tawil et al., 2003; Read,
1985) which create an overall shallow water depositional system that is susceptible to subaerial exposure during relative low stand in sea-level. Since the vulnerability to the effects of dissolution is high in carbonate rocks, subaerial exposure can diagenetically alter the original material creating distinct geochemical signatures (Allen and Matthews, 1982). As such, surfaces of subaerial exposure can be identified in carbonates geochemically where weathering profiles are not well developed.
Meteoric diagenesis of subaerially exposed carbonate rock can be identified by excursions patterns in δ13C and δ18O values and trace elements (e.g., Sr) across stratigraphic horizons. Photosynthesizers preferentially incorporate 12C during the production of biomass; this
carbon may return to the environment as soil gas CO2 during respiration or oxidation of organic
matter (Allan, 1976; Allen and Matthews, 1982; Beier, 1987; Goldstein, 1991; Meyers and
Lohmann, 1985). Meteoric water that infiltrates the soil horizon may pick up this CO2 and
become acidic, leading to dissolution during the descent of this water to the water table.
Subsequent recrystallization incorporates the isotopically light soil-gas derived carbon into
calcite precipitates (Allan, 1976; Allan et al., 1977; Railsback et al., 2003; Theiling et al., 2007)
providing a marker of this diagenetic event. 89
Meteoric water derived from the evaporation of seawater is isotopically light with respect to oxygen. As sea water evaporates, water molecules containing 16O go into the gas phase more
readily than heavier 18O containing water molecules. This results in meteoric water being
isotopically lighter relative to sea water where carbonate sediments are being generated
physically and biogenetically. Isotopically light oxygen in meteoric solution infiltrating marine
derived carbonate sediment at subaerial exposure surfaces can combine with soil-gas CO2 to
form an acidic solution which dissolves carbonate. If the infiltrating meteoric solution becomes
saturated with respect to calcite, carbonate precipitates can form with lighter meteorically
derived oxygen incorporated into calcite precipitate during recrystallization resulting in a
diagenetic marker of the event. Carbonate precipitates associated with subaerial exposure
surfaces exposed to humid conditions will display a negative excursion marker in δ18O relative to marine derived carbonate. However, in semi-arid/arid environments the excursion marker in δ18O is positive due to evaporative fractionation of the lighter 16O water molecule at the air-sediment
interface.
The trace element Sr is a common constituent of marine derived aragonite. Meteoric
water infiltrating marine sediment at subaerial exposure surfaces readily dissolves aragonite
sediment. The Sr released from the marine rock during dissolution does not incorporate
appreciatively back into calcite precipitates and most is lost from the rock system as the meteoric
water continues to move through the vadose zone (Theiling et al., 2007). The preferential
exclusion of Sr in calcite precipitates produces a diagenetic marker characterized by a net
decrease in weight % Sr relative to the original marine carbonate rock.
Many studies have documented the meteoric diagenesis of carbonate rocks and sediments
using stable isotopes δ13C, δ18O and bulk trace elements. Pioneering work in carbonate 90 diagenesis delineated isotope patterns within the meteoric environment and provided a model to interpret Pleistocene limestones (Allan et al., 1977; Beier, 1987; Benson and Matthews, 1971;
Land, 1970; Land and Epstein, 1970; Videtich and Matthews, 1980). Further use of the meteoric diagenesis model has been extended to include the diagenesis of Paleozoic carbonate rocks
(Algeo et al., 1992; Allen and Matthews, 1982; Goldstein, 1991; Meyers and Lohmann, 1985;
Wu et al., 2005), as well as porosity preservation in economically important limestone reservoir rock (Humphrey et al., 1986; Wagner and Matthews, 1982), terrestrial plant evolution (Beeunas,
1984; Keller and Wood, 1993; Watanabe et al., 2000) and sequence stratigraphic boundary delineation (Railsback et al., 2003; Theiling et al., 2007).
Previous studies have identified surfaces of subaerial exposure by sampling rocks with along a single vertical transect (Allen and Matthews, 1982; Beier, 1987; Goldstein, 1991;
Railsback et al., 2003). Workers sampling along a single vertical transect assumed that intrastratal lateral geochemical variation was sufficiently small to preclude the incorrect assignment of an exposure surface. However, Theiling et al., (2007) recognized potential errors that could arise when sampling along a single vertical transect. Heterogeneity in the distribution of biotic cover, porosity and permeability variability in exposed sediment, and differences in carbonate mineralogy along a syndepositional stratum can substantially affect meteoric diagenesis leading to intrastratal geochemical heterogeneity. For instance, replicate sampling (10 samples per horizon collected 2 meters apart over a horizontal distance and separated 10 centimeters vertically from each horizon) along stratigraphic horizons above and below the M4-
M5 sequence boundary of Holland and Patzkowsky (1997; 1998) in an outcrop of Ordovician limestones near Gladeville, Tennessee, revealed considerable variance in δ13C and δ18O values
and Sr content among seemingly identical samples (Theiling et al., 2007). These results suggest 91 variance in geochemical expressions can potentially contribute to the identification of subaerial exposure surfaces. Furthermore, Theiling et al. (2007) showed that replicate sampling allowed statistical arguments to be made for the presence of subaerial exposure that were previously undetected. Although Theiling et al. (2007) delineated geochemical heterogeneity along a subaerial exposure surfaces, they did not evaluate the scale of heterogeneity that exists along a stratigraphic horizon. In short, can geochemical heterogeneity observed by Theiling et al. (2003) be recognized in hand sample or is it only observable in sampling schemes where samples are collected meters apart along a single horizon? If geochemical heterogeneity exists at the scale of a hand sample then the replicate sample methodology described by Theiling et al., (2007) could be applied to the study of cores.
Many studies have demonstrated the utility of mineral characterization in reconstructing basin development within a sequence stratigraphic framework (Al-Ramadan et al., 2005; Burns et al., 2005; El-ghali et al., 2006; McConachie and Dunster, 1996; Morad et al., 2000; Taylor et al., 2000; Taylor et al., 1995). Syndepositional dolomite deposited in tidal flat facies could serve as an important marker for sequence definition within a sequence stratigraphic framework.
Dolomite is currently forming as thin beds in many evaporitic tidal flats on the southern and western margins of the Arabian Gulf (Illing and Wells, 1964; Mackenzie, 1981; Warren, 2000;
Wenk et al., 1993) which is the modern analog for much of the Mississippian aged rocks in the
Appalachian Basin. For example, Swart et al. (1987) observed well-formed dolomite crystals precipitating instantaneously in the hypersaline water in a pit dug in coarse quartz sabkha sand of the Qatar Peninsula. Fine crystals of dolomite have been observed growing around and partially replacing aragonite needles in pore space within crusts buried beneath an Abu Dhabi sabkha
(Mackenzie, 1981). Very finely crystalline dolomite in tidal flat facies may represent 92 penecontemporaneous sabkha-like dolomite formation (Scholle and Ulmer-Scholle, 2003). Al- tawil (1997) visually recognized the presence of dolomite in outcrop below surfaces of subaerial exposure in the Mississippian rocks of the Appalachian Basin; however, no attempts were made to quantify the minerals concentration. Petrographic analysis of this dolomite reveals very finely crystalline dolomite suggesting a primary origin and not a later burial process. Quantitative XRD of bulk rock could prove useful in the characterization of dolomite in association with subaerial exposure surfaces.
Studies have demonstrated that the development of a predictive model of clay mineral distribution patterns in sedimentary sequences within a sequence stratigraphic framework is possible (Al-Ramadan et al., 2005; Amorosi, 1995; Amorosi and Centineo, 1997; Ketzer et al.,
2003a; Ketzer et al., 2003b; Salem et al., 2005; Środen, 1999). These models are dependent on the idea that the type and abundance of clay minerals formed along subaerially exposed sequence boundaries are strongly climate dependent (Ketzer et al., 2003a). Mg-rich smectites, precursor minerals to corrensite, can form authigenically in modern semiarid/arid tidal flat environments
(Hover et al., 1999; Ketzer et al., 2003a; Martini et al., 2002). Corrensite, a regularily ordered
(50:50) interstratified chlorite-smectite clay mineral shown to be an intermediate phase in the
Mg-smectite-corrensite-chlorite transition, is the product of burial diagenesis of Mg-rich smectites. In oriented clay mounts, corrensite is easily identified in X-ray diffraction (XRD) patterns by a strong reflection near 26.5-29 Å in air dried state and a shift to 30-31Å after glycolation (Figure 4.1). Many studies have described corrensite in association with ancient marine evaporites and tidal flat environments (Andreason, 1992; Bayhan, 2007; Bodine, 1985;
Fisher, 1988; Han et al., 2000; Hillier, 1993; Jeans et al., 2005; Rao, 1992; Ryan and Hillier,
2002; Stein et al., 1991). Since many of the bounding surfaces in the study area are capped by 93
Figure 4.1: XRD pattern for a regularly ordered (R=1) 50:50 interstratified trioctahedral chlorite and trioctahedral smectite (corrensite) generated with the program Newmod: A) Air-dried state; B) ethylene glycol-solvated. 94 restricted marine tidal flat deposits (Al-Tawil and Read, 2003), the presence of corrrensite in these rocks provide a marker for deposits associated with restricted hypersaline tidal flat environments and subaerially exposed surfaces formed during semiarid-arid climate.
In light of these issues, the present study has applied the replicate sample method to hand-sized samples taken along a single vertical transect across a Mississippian sequence boundary independently defined by Al-Tawil and Read (2003), which exhibits macroscopic and microscopic features of subaerial exposure, to determine if heterogeneity in geochemical expression exists at a smaller lateral extent than that described by Theiling et al., (2007). Specific questions addressed in this research include: 1) do geochemical signatures of subaerial exposure vary significantly at the scale of a hand samples; 2) can statistical analysis of replicate sampling from hand samples provide meaningful information concerning the position of surfaces of subaerial expsosure; 3) can whole rock mineralogy and clay mineralogy provide additional constraints on the recognition of subaerial exposure surfaces?
Regional Background (Sequence Stratigraphy)
High-resolution sequence stratigraphic analysis of Mississippian strata (Meramecian to
Chesterian) in Kentucky records the development of thirteen, 5 to 20 m thick, fourth-order depositional sequences (Al-Tawil and Read, 2003). Although relatively thin on the distal foreland, these fourth order sequences thicken significantly to the east and northeast into the proximal foreland region of Virginia and West Virginia, where they attain thicknesses of 30 to
90 meters (Al-Tawil et al., 2003). The sequences are stacked into several siliciclastic-bounded third-order composite sequences that make up a broadly transgressive Mississippian supersequence. The M1-C1, C4-C5 and C8-C9 boundaries of (Al-Tawil and Read, 2003) represent well-developed boundaries for these third-order composite sequences. The sequences 95 were deposited during rapid sea-level rise coupled with increased differential subsidence and they represent high stand systems tracts where the surface of maximum flooding coincides with the transgressive surface. In contrast, sequences developed during transgression that did not greatly exceed sedimentation rates are composed of transgressive and highstand systems tracts.
The C8-C9 sequence boundary of Chesterian age marks a significant transition in depositon in the Appalachian Basin. All sequences up to C8 were deposited during relatively moderate amplitude sea-level fluctuation that flooded the ramp to depths of 10 meters or less
(Al-Tawil and Read, 2003). Arid to semiarid conditions coupled with moderate sea-level flucuation resulted in relatively thin sequences dominated by ooid grainstone and partitioned by caliches, lime mudstone and eolinite at the sequence boundaries. Sequences after C8 were deposited during sea-level changes of greater magnitude, most likely representing tens of meters in amplitude (Al-Tawil and Read, 2003). The increase in the magnitude of sea-level change was most likely the result of Milankovitch long-term eccentricity forcing during the onset of global icehouse conditions on Gondwana (Al-Tawil and Read, 2003; Al-Tawil et al., 2003; Smith et al.,
2000; Smith and Read, 2001; Wynn and Read, 2006). These later sequences, which were formed under more humid conditions, are dominated by widespread skeletal grainstones, wackestones and they are partitioned by clay-rich paleosols and marine siliciclastic units.
The Pine Mountain Outcrop
A continuous well-exposed five hundred foot section of the Mississippian system (Figure
4.2) is exposed in a road cut about 4 km long where I-75 climbs Pine Mountain on the Pine
Cumberland overthrust sheet just south of Jellico, Campbell County, Tennessee. Sequences M1 and C1 through C12 are well exposed and easily accessible at this location. The Reelsville at
Pine Mountain, a five meter thick unit in the upper C8 sequence, is a relatively complex 96
Figure 4.2: Map of the outcrop locality on the Pine Cumberland overthrust sheet on I-75 southeast of Jellico, Campbell County, Tennessee.
96 97
Figure 4.3: Stratigraphic column of the C8 sequence exposed on the Pine Cumberland overthrust sheet on I-75 southeast of Jellico, Campbell County, Tennessee.
98 unconformity bounded unit deposited during a significant sea-level fall associated a third-order super sequence boundary (Al-Tawil and Read, 2003). The entire Reelsville unit of the C8 sequence and the lowermost two meters of the C9 sequence is the focus of this study. The stratigraphic positions reported in this paper are based on the total Pine Mountain section reported in Al-tawil (1997).
This section of the C8 sequence lies within the Beaver Bend Formation which transitions upward into the 4.5 m disconformity bounded Reelsville unit (Figure 4.3). Thlower 7 m section is composed exclusively of oolitic grainstone. An abrupt facies transition separates the oolitic section from the skeletal grainstone-wackestone-mudstone unit of the Reelsville. Traditional facies analysis of the Reelsville unit reveals three fining upward parasequence cycles composed of skeletal grainstone –wackestone and dolomitized mudstone. Facies relationships are complex in the Reelsville, however, petrographic analysis of the unit suggests it was deposited under limited accommodation space such as that imposed by inner ramp lagoon and peritidal mudflat environments. A 30 cm breccia horizon, interpreted as the C8-C9 sequence boundary, separates the Reelsville unit from the overlying Beech Creek Formation (Figure 4.4). The C9 sequence comprises the entire Beech Creek Formation. It is approximately 11 m thick and composed mostly of skeletal grainstone-packstone; it is composed mostly of echinoderm and brachiopod fragments suggesting deposition under normal marine salinity conditions.
Physical Evidence for Lowstand Deposits and Subaerial Exposure
A depositional sequence is a relatively conformable succession of genetically related strata bounded at the base and top by unconformities and their correlative conformities (Van
Wagoner, 1995). Unconformities are generated during sea-level lowstands when upper bounding surfaces of sediment or rock are exposed at the Earth’s surface. During exposure, 99
Figure 4.4: The C8/C9 sequence boundary: A) the contact between the upper Reelsville unit (upper C8) and overlying Beech Creek Formation (lower C9). The C8/C9 sequence boundary is highlighted with a red line. B) Collapse breccia immediately below sequence the boundary. 100 subaerial diagenetic processes modify or obliterate pre-existing fabrics. This process typically results in a break in the sedimentary record. Recognition of these surfaces is dependent upon the total time of exposure, extent of modification in the rock, the extent that the subaerial diagenetic horizon is preserved in the rock record and our powers of resolution and observation (Esteban and Klappa, 1983). By definition of a depositional sequence one can see quite quickly that identification of subaerial exposure surfaces is a fundamental task in sequence stratigraphy; however, this is often difficult in carbonate rock successions due to the shallow nature of deposition of these systems and the ease at which dissolution and removal of dissolved material can occur during subaerial exposure. In carbonate systems this can often result in a cryptic surface of subaerial exposure (Theiling et al., 2007).
Diagenetic alteration of carbonate sediment or rock can lead to recognizable features at subaerial exposure surfaces. The type of structure preserved at subaerial exposure surfaces can differ depending upon climate conditions and exposure time. For example, surfaces of subaerial exposure in carbonate rock developed over relatively long periods of time under humid climatic conditions would favor the development of a siliciclastic-rich paleosol horizon whereas under hyperarid-arid climate conditions, a caliche horizon would develop. These two end member structures are very different diagnostically, which highlights the importance in having a set of descriptive criteria when identifying subaerial surface exposures that spans between humid and arid regimes. To this end, the following is a characterization of features used to identify subaerial exposure surfaces in carbonates during periods of humid and arid climate regimes.
Structures Associated with Subaerial Exposure
Karst is a diagenetic overprint in subaerially exposed carbonate rock produced by differential dissolution and transport of dissolved materials by meteoric water. In outcrop, karst 101 surfaces can be recognized by an irregular surface that truncates grain and bedding structures
(Figure 4.5A). Petrographic analysis of grainstones and packstones below karst truncation features reveal grain dissolution (Figure 4.5B). Sometimes a breccia horizon, the result of the collapse of material into a dissolution cavern, develops along the surface (Figure 4.5C). In a few cases updip where karst is well developed, the collapse feature can be quite large and easily recognizable by the alteration of bedding above the collapse feature (Figure 4.5D).
Siliciclastic-rich paleosols are well developed on surfaces exposed for relatively long periods of time under humid conditions (Al-Tawil and Read, 2003). These deposits are fine- grained siliciclastic-rich units with orange-red-green mottling (Figure 4.5E). Petrographic inspection often reveals highly degraded carbonate skeletal grains from the original carbonate material (Figure 4.5F). Mottling is the result of differences in redox conditions within the paleosol during its formation. In many cases root traces and minor coal traces can be observed.
Transgressive lags appear as conglomeritic to breccia-rich interclast/intraclast horizons immediately above a subaerial exposure surface (Figure 4.5G). The lag is produced during the reworking of previously exposed material by tides during the onset of sea-level rise and therefore represents a marker bed above surfaces of subaerial exposure. Some lags are composed of well- rounded clasts while others appear angular to subangular in shape. In the study area, the lag is typically restricted to a one foot interval immediately above an exposure surface. However, lags do not always appear above all subaerial exposure surfaces in the study interval.
Tidal flat facies is composed of massive fine-grained dolomite and siliciclastic material.
Fossils are typically rare to completely absent (Figure 4.5H). When present, fossils are represented by small disarticulated fragments of echinoderm and ostracod material. Quartz pseudomorphs after gypsum appear locally in some tidal flat facies in horizons beneath 102 truncation features interpreted as surfaces of subaerial exposure. Petrographic analysis of tidal flat facies sometime reveal calcite-filled desiccation cracks exhibiting geopetal fill features. In outcrop, tidal flat faces stand out by their massive fine-grained texture and darker color, which exhibits high contrast with the lighter colored calcite limestone.
Restricted marine facies is composed of massive to clotted fine-grained micrite. Fossils are represented by articulated and disarticulated ostracod shells with minor echinoderm ossicles
(Figure 4.5I). The depositional environments for restricted marine facies in the study area are restricted lagoon or intershoal settings. Restricted marine facies will sometimes display truncation features indicative of periods of subaerial exposure. Quartz pseudomorphs after microbial mat/stromatalite and mudcracks have been observed in this particular facies.
Quartz pseudomorphs after gysum appear as cauliflower shaped microcrystalline nodules typically centimeters to tens of centimeters below surfaces of subaerial exposure within the tidal flat facies (Figure 4.5J). Quartz pseudomorphs after gypsum nodules are associated with dolomite-rich tidal flat facies. Petrographic inspection often reveals small inclusions of anhydrite, presence of length slow quartz (Figure 4.5K), zebraic quartz (Figure 4.5L) and/or fortification zoning (Figure 4.5M). The formation of length slow and zebraic quartz and its association with the replacement of evaporite minerals has been well documented (Chowns and
Elkins, 1974; El Khoriby, 2005; Elorza and Rodriguez-Lazaro, 1984; Folk and Pittman, 1971;
Gomez-Alday et al., 2002; Maliva, 1987; McBride and Folk, 1977; Milliken, 1979; Overstreet et al., 2003; Siedlecka, 1976; Swennen et al., 1990; Tucker, 1976; Ulmer-Scholle and Scholle,
1994; Ulmer-Scholle et al., 1993)
Quartz peudomorphs after microbial mat/stromatalite colonies appear as rounded concentric layered nodules in association with shallow restricted marine facies, truncations and 103 dessication features (Figure 4.5N). Many of the nodules have collapse and geopetal fill features suggesting silicification occurred prior to sediment consolidation. Petrographic analysis reveals fine-grained inclusions within the microcrystalline quartz that take on a layered orientation.
These inclusions may have been sediment trapped in the outermost layer of the mat during growth.
Fenestral fabrics consist of a series of irregular to lamination parallel cavities that are typically larger than sediment grains (Figure 4.5O). In some examples the cavities contain geopetal fill. Fenestrae are generally associated with microbial mats and result from the shrinkage, gas formation, organic decay, trapping of air through swash-zone wave action, or other synsedimentary processes (Choquette and Pray, 1970). Fenestral fabrics are of particular importance because they commonly denote peritidal deposition (Scholle and Ulmer-Scholle,
2003).
Eolinites are wind-driven subaerial accumulations of carbonate-dominated and carbonate- cemented sand (Brooke, 2001). Wind generated deposits are heterogeneous in size and often coarse-grained, making them prone to misinterpretation as a marine deposit (Frebourg et al.,
2008). Eolinites are composed of well-rounded and abraded peloids, whole and abraded ooids, skeletal grains and finer-grained quartz. In outcrop, eolinite deposits exhibit low angle to nearly planar bedding and a pin-striped appearance. Petrographic analysis reveals grains that are well- sorted and tightly packed within laminae and exhibit reverse grading (Figure 4.5P). Many individual grains appear abraded and some ooids may show fracturing and subsequent rounding.
Red beds appear as tan-orange to red-orange or maroon beds of fine-grained dolomite and fine-grained siliciclastic material (Figure 4.5Q). The oxidized sediment sometimes appears detrital based upon its orientation relative to bedding. Other beds appear as post depositional 104 insitu oxidation of the sediment. Fossils are rarely found if at all in red beds. (Al-Tawil and
Read, 2003) interpret red beds within the study interval as marginal marine settings subject to periodic exposure.
Rhizoliths are organosedimentary structures produced in roots by accumulation and/or cementation around, cementation within, or replacement of, higher plant roots by mineral matter
(Klappa, 1980). Typically they are calcite filled, vertically oriented bifurcating structures whose diameter decreases with penetration depth (Figure 4.5R). When observed in the field or thin section, rhizoliths represent definitive evidence for
Desiccation cracks are vertical to subhorizontal cracks that are the result of contraction and shrinkage of wetted mud during desiccation in the tidal flat environment. In plan view, mudcrack polygons are a typical macroform of the desiccation process; however, in outcrop this view is not always available. Petrographic analysis often reveal calcite filled microfractures in tidal flat facies (Figure 4.5S). The fractures sometimes give the clotted micrite fabric
(characteristic of tidal flat facies) an incipient crumbly appearance. Desiccation cracks sometimes form circumgranular patterns around skeletal grains or indurated peds. In many cases the cracks show evidence of minor sediment collapse and geopetal fill.
Caliche appears as fine-grained, chalky to well-cemented calcite with a variable massive to irregular platy, subhorizontal crust laminae. It is typically light in color making it easily recognizable in outcrop. Petrographic analysis of sheet-like caliche suggests the predominant fabric is clotted-peloidal micrite (Figure 4.5T). Massive caliche is characterized by poorly cemented white to cream calcite or dolomite with a chalky consistency.
105
Caliche is a pedogenic feature that cap and therefore define subaerial exposure surfaces associated with hyper-arid to arid conditions (Al-Tawil and Read, 2003; Esteban and Klappa,
1983; Ettensohn et al., 1988).
Methods
Twenty three hand samples were collected perpendicular to the outcrop face from the 8 meter
C8-C9 boundary section. A total of 20 samples were collected below the C8-C9 boundary and 6 above. Samples were collected along a single vertical transect approximately every 30 cm.
Hand samples were slabbed with a water-cooled tile saw, polished and sonified for one minute to remove any fines adhering to the surface. The slab surface was gridded into ten ~4 cm2 areas. A dental drill was used to isolate micrite ~300 µg
from the surface of each area (Figure 4.6). Drill speeds were kept low to avoid alteration effects
related to high speed drilling (Gill et al., 1995).
Thin sections for each sample were prepared with Hillquist thin sectioning equipment at
the University of Georgia. Thin sections were used for classification of samples according to
Dunham (1962) and facies confirmation. Conventional facies analysis was performed based on
the descriptions in Al-Tawil and Read (2003) and Al-Tawil et al. (2003), which is consistent
with contemporaneous facies analysis described.
elsewhere in North America (Carney and Smosna, 1989; Ettensohn et al., 1984; Leonard, 1968;
Smith et al., 2000; Smith and Read, 2001; Wynn and Read, 2006).
Quartz, calcite and dolomite were quantitatively assessed by XRD. Bulk rock samples
were first lightly crushed with a hammer. Approximately 8 g of the less than 2mm fraction 106
Figure 4.5: Physical features of subaerial exposure; A) Two irregular truncations in strata (subaerial exposure surfaces) at the C4/C5 sequence boundary at Jellico, Tennessee; B) Photomicrograph showing grain dissolution associated with subaerial exposure; C) Breccia horizon associated with the C8/C9 sequence boundary at Jellico, Tennessee; D) Solution hole collapse feature at Moorehead, Kentucky; E) Red and green mottled, fine-grained siliciclastic- rich paleosol at Moorehead, Kentucky; F) Photomicrograph of a highly degraded allochem (possible brachiopod fragment) in fine-grained siliciclastic-rich paleosol.
107
Figure 4.5 continued: Physical features of subaerial exposure; G) transgressive lag in hand sample (black arrows pointing toward intraclasts-interclasts) collected above C9/C10 sequence boundary (scale divided into 1 cm increments); H) photomicrograph of fine-grained dolomite rhombs and siliciclastic material characteristic of a tidal flat environment; I) photomicrograph of articulated and disarticulated ostracod with minor echinoderm elements in clotted fine-grained micrite characteristic of a restricted marine environment; J) quartz pseudomorphs after gypsum nodules (outlined in red) in association with dolomite-rich tidal flat facies; K) photomicrograph of length-slow quartz in a quartz nodule interpreted as a pseudomorph after gypsum under cross polarization and insertion of the length-fast gypsum plate from the SE. The colors in the NW and SE quadrants decreased, which means that it is length-slow; L) photomicrograph of zebraic quartz in quartz nodule interpreted as a pseudomorph after gypsum. 108
Figure 4.5 continued: Physical features of subaerial exposure; M) photomicrograph of cubic terminations in chalcedony (fortification zoning) interpreted as replacement of evaporite minerals such as halite; N) quartz peudomorphs after microbial mat/stromatalite colonies appear as rounded concentric layered nodules (red arrow) in association with shallow restricted marine facies. Note truncation in strata (top of rock hammer) and numerous mud cracks radiating from nodules; O) photomicrograph of fenestral cavity with geopetal fill; P) photomicrograph showing reverse grading in well-rounded tightly packed carbonate sediment characteristic of eolinite deposits; Q) maroon colored beds of fine-grained dolomite and siliciclastic material interpreted as marginal marine in origin (scale divided into 1 cm increments); R) photomicrograph showing vertically oriented bifurcating structures whose diameter decreases with penetration depth representing rhizoliths.
109
Figure 4.5 continued: Physical features of subaerial exposure; S) photomicrograph of calcite filled microfractures in tidal flat deposit; T) photomicrograph of clotted-peloidal sheet-like caliche.
110
Figure 4.6: Representative example of the subsample collection scheme for each hand sample. Slabbed surface is gridded into ten ~4 cm2 areas. Dental drill was used to isolate micrite (~300µg) from the surface of each area (red circles).
111 was placed in a McCrone Mill with methanol and ground for 10 minutes. Randomly oriented bulk powders were analyzed with a Bruker D8 Advance XRD using the following parameters: 2 to 75o 2 Θ, 40 ma, 40 kv at 5o per min step and 0.02 increment. Quartz, calcite and dolomite
percentage were calculated with a Rietveld refinement algorithm provided in the program Topas
® by Bruker. Precision of the Rietveld refinement was assessed with known mixtures of quartz,
dolomite and calcite standards and found to be within ± 3%.
Oriented clay slides were analyzed with XRD for clay mineral identification.
Approximately 20 grams of crushed rock was allowed to react with a 15% solution of acetic acid for 24 hours. The separated constituents were repeatedly washed in deionized water, disaggregated in a solution of deionized water and hexametaphosphate and the <2 µm fraction separated by centrifugation. Clays were concentrated by allowing the solution containg the 2 µm fraction to evaporate, resaturating the dried clay fraction with 5 ml of deionized water, sonified for approximately 10 seconds and pipetted onto a petrographic slide. The dried slides were analyzed with a Bruker D8 Advance XRD using the following parameters: 2 to 35o 2 Θ. 40 ma,
40kv at 5o per minute step and a 0.02 increment. Clay slides were run both dry and ethylene
glycol saturated.
Stable isotope analysis of carbon and oxygen was conducted using a Gasbench II device
coupled to a Delta Plus XP isotope ratio mass spectrometer configured in continuous flow mode.
CO2 gas was extracted from powdered samples using the phosphoric acid digestion method
(Craig, 1957). Samples were reacted at 50 oC. International reference standard material used was
NBS-19 limestone. Samples were calibrated to VPDB through the repeated analysis of NBS-19
every sixth sample throughout a sequence. Precision was determined by replicate analysis of 59
samples and are 0.15 for both δ13C and δ18O. Ten sub-samples from each hand sample were 112 analyzed separately for δ13C and δ18O. The δ13C and δ18O sub-sample values were collectively assessed to determine the arithmetic mean:
1 n x = • ∑ xi n i
where x = arithmetic mean, n = the number of terms and xi =the value of each individual item in
the list of numbers being averaged for each stratigraphic horizon. In addition, the values were
collectively assessed to determine the variance:
1 n S 2 = • ∑ ( x - xi )2 n i
where S2 = variance, n = the number of terms, x = the arithmetic mean and xi = the value of
each individual item in the list of numbers for each stratigraphic horizon.
Results
A parasequence is a relatively conformable succession of genetically related beds or
bedsets bounded below and above by flooding surfaces and their correlative surfaces (Van
Wagoner, 1995). The upper C8 sequence contains three parasequences within the Reelsville unit
(Figures 4.7, 4.8 and 4.9) based upon the resolution afforded by the petrographic examination of
hand samples. All three parasequences show a shallowing upward trend that begins with either a
normal marine or restricted marine bed set that shallowed to a tidal flat. In some cases evidence
for subaerial exposure was observed in thin section which could be correlated with field
observations, while others did not. Field observations were combined with petrographic
observations to accurately define apparent subaerial exposure surfaces. 113
113 Figure 4.7: Facies interpretation of Parasequence 1 of the Reelsville unit (upper C8 sequence) at the Pine Mountain locality. Scale bar in the photomicrographs is equal to 1 mm. 114
Figure 4.8: Facies interpretation of Parasequence 2 of the Reelsville unit (upper C8 sequence) at the Pine Mountain locality. Scale bar in the photomicrographs is equal to1 mm.
114
115 115 the Pine Mountain locality. S cale bar ille unit (upper C8 sequence) at of Parasequence 3 the Reelsv Figure 4.9: Facies interpretation is equal to1 mm. in the photomicrographs 116
The base of the Reelsville is represented by a dolomitized pelletal mudstone, displaying clotted fabric and containing minor ostracod-mollusk fragments and root traces. The underlying surface is a truncation of the Beaver Bend Formation, which is composed almost exclusively of ooid grainstone. The base of the first parasequence (P1) is a well sorted, fine-grained peloidal brachiopod-ostracod packstone that fines upward into a mudstone containing ostracod elements and finally a mudstone absent of fossils (Figure 4.7). The base of the second parasequence (P2) is a skeletal wackestone that fines upward into an ostracod wackestone and finally a mudstone completely devoid of fossils (Figure 4.8). The base of the third parasequence (P3) is an ostracod- spiculiferous wackestone exhibiting fenestrae with geopetal structure that passes upward into a peloid-intraclast packstone and finally an ostracod and pelletal dolomite-rich mudstone penetrated by rhizoliths (Figure 4.9). Dolomite present in the mudstone samples of all three parasequences appear as fine-grained rhombs within the mud matrix.
The Reelsville is a depositionally complex, well-developed unconformity-bounded unit that is the result of a significant sea-level fall associated with a third-order super sequence (Al-
Tawil and Read, 2003). Seven physical features were observed in the Reelsville unit during field observation that represents potential regions of subaerial exposure. Figure 4.10 shows the seven physical features in relation to the three parasequences comprising the Reelsville unit. A summary of their characteristics are summarized in Table 4.1. Field interpretation of the
Reelsville unit was difficult due to the complexity in sedimentation patterns, the overall fine- grained nature of the deposits and the minimal development of pedogenic features due to the overall arid climate. Although several truncation features were observed, it was difficult to 117
Figure 4.10: Physical evidence for subaerial exposure observed in the Reelsville unit. Numbers 1 through 7 correspond to seven features identified during field observations.
118
Table 4.1: Characteristics and stratigraphic position of seven physical features indicative of subaerial exposure observed in the Reelsville unit during field reconnaissance. 142
118 119 distinguish subaqeuous from subaerial erosional features by field observation alone. Petrographic observation became necessary to accurately classify samples, construct parasequence patterns and identify small sedimentary structures indicative of subaerial exposure. In light of the issues described above, geochemical proxies became imperative for definitive identification of subaerial exposure surfaces.
Whole rock analysis of powders with XRD reveals differences in quartz, dolomite and calcite concentration in the upper C8 and lower C9 sequence. For the Reelsville unit, the range in dolomite concentration is between 5 and 89 %. Dolomite concentrations are highest at 46.3 (89
%), 46.8 (19%), 48.4 (44%) and 49.3-51.1 (64-84%) meters above base of section (Figure 4.11).
Overall, there is an increase in dolomite concentration from the surface of subaerial exposure at the base of the Reelsville unit towards the C8-C9 sequence boundary. In contrast, the lower 2 meters of the C9 sequence, dolomite range is approximately 3 to 4%. The C9 sequence is dominated by calcite which averages 90 %. One surprising aspect of the data is that dolomite concentration is highest below apparent surfaces of subaerial exposure previously identified during field reconnaissance (Figure 4.10) with the exception of the one at the 47.4 m position
(#3). This trend in concentration is not recognizable in the field or petrographic observations due to the fine-grained nature of the allochems and matrix in the samples.
The trend in quartz concentration is less definitive for identifying subaerial exposure surfaces.
For the Reelsville unit, the range in quartz concentration is between 4 and 19%. Two peaks in quartz concentration correlate with apparent subaerial exposure features #4 at 48.7-48.8 m (19%) and #6 at 50.1- 50.6 m (18%). Overall, there is a slightdifference in quartz concentration above 120
Figure 4.11: Bulk mineralogy, clay mineralogy, mean δ 13C and δ 18O, δ 13C and δ 18O variance of samples collected from the Pine Mountain locality. Solid black = quartz, grey = dolomite, yellow = calcite and orange = corrensite.
120 121 and below the C8-C9 sequence boundary. The range in quartz concentration of the two meter interval below the C8-C9 is between 6 and 19% whereas the above it is between 6 and 9%.
The clay mineral corrensite is a ubiquitous component of the clay suite of the Reelsville unit. It first appears 46.8 meters above base of section, which is immediately below the first appearance of restricted marine facies (ostracod wackestone) in the P1 parasequence (Figure
4.7), and continues up to the C8/C9 boundary (Figure 4.11). Corrensite is not present immediately above the sequence boundary (lower C9 sequence) nor is it seen in the C8 sequence below 46.8 meters. These strata are normal marine deposits that have not been influenced by hypersaline reflux. The clay minerals in these deposits are mixed-layered illite-smectite, chlorite, illite and minor kaolinite (Figure 4.12). The appearance of corrensite in strata of all three parasequences with diverse fauna, indicative of normal marine conditions, suggests that the reflux of hypersaline water through these sediments occurred during the deposition of restricted marine and tidal flat facies.
There are three major negative excursions in mean δ13C values within the Reelsville unit
(Figure 4.11). The excursions occur between 45.4 and 46.3 meters, 49.3 and 50.2 meters and
50.2 and 51.3 meters. After the first excursion between 45.4 and 46.3 meters, there is a steady
decrease in mean δ13C value up to the second major excursion at 49.3 and 50.2 meters. The two
upper excursions are separated by a relative high in mean δ13C value suggesting that they represent two separate subaerial exposure events. Above the C8/C9 sequence boundary, which is represented by the trangessive surface and a return to normal marine conditions, there is a high invariant mean δ13C value. Overall, the mean value of δ13C for the entire Reelsville unit is 0.02
‰ VPDB compared to a mean value of 1.71 ‰ VPDB for the 0.9 meter horizon above
122
Figure 4.12: Diffractogram of the clay mineral fraction (<2µm) from a sample interpreted as normal marine (ooid grainstone) from the Beaver Bend Formation.
123 the C8/C9 sequence boundary (lower Beech Creek Formation). In addition, negative excursions maxima in mean δ13C values correspond with increases in dolomite concentrations measured in
bulk rock XRD analysis. Excursion maxima and minima in mean δ13C values were assessed for
statistical significance with Student’s t-test (Figure 4.13). Student’s t- test results across
excursion horizons presented in Table 4.2 suggest that all three excursion patterns in δ13C values
represent statistically significant differences in mean.
There are four relative highs in δ13C variance within the Reelsville unit (Figure 4.11).
Relative highs occur at 46.3 and 47.8, an interval between 48.5 to 50.7 and 51 meters above base
of section. The relative high at 47.8 does not appear to correspond with a negative excursion in
mean δ13C value but the remaining three do appear to correspond to the major excursions in mean δ13C value described previously described. Horizons of high variance were compared to
horizons of low variance immediately above (Figure 4.13) were assessed for statistical
significance with an F test (Table 4.2). F test results suggest all four horizons are statistically
significant. Similarily, regions of high variance correspond closely with high dolomite
concentrations.
The Reelsville unit is characterized by an overall positive excursion trend in mean δ18O value. The largest excursion in the mean value of δ18O occurs at the surface of subaerial
exposure at the base of the Reelsville unit (46.3 m), which is a 4 ‰ difference in mean from the
sample immediately below the surface. A positive trend exists between 46.5 to 50.5 m and then a
sharp decrease thereafter. Within the broad positive trend, there are five small positive
excursions, which are 0.5 to 2 ‰ in magnitude (Figure 4.11). These smaller positive excursions
may represent surfaces of subaerial exposure within the Reelsville unit. Several of the
124
Figure 4.13: Statistical analysis of δ13C values. The green box is the upper horizon and the orange box the lower horizon of values analyzed for each surface of subaerial exposure.
125
Table 4.2: Student’s t-test of difference of mean and F test of difference of variance results of δ13C and δ18O for Horizon 1 through Horizon 6.
126 excursions in the mean δ18O value correspond with excursions in dolomite %, mean δ13C and
δ13C variance. Excursion maxima and minima in mean δ18O values were assessed for statistical
significance with Student’s t- test (Figure 4.14). Student’s t-test results across excursion horizons
with the horizon immediately above suggest that all six horizons are statistically significant
(Table 4.2).
Three major and two minor relative highs in δ18O variance exist within the Reelsville unit
(Figure 4.11). These highs correspond with five of the six positive excursions in the mean δ18O value. The one exception is at position 50.5 m. This area corresponds with a significant negative
excursion in the mean of δ13C samples, high variance in δ13C samples and a major positive
excursion in the mean of δ18O samples. Relative highs in δ18O variance correspond with relative
highs in dolomite %. F-test results across high variance horizons in δ18O of samples below
surfaces of exposure (Table 4.2) suggest five of the six horizons are statistically significant.
Collectively, stable carbon and oxygen isotope data, dolomite % and the occurrence of the clay
mineral corrensite show patterns that correlate with petrographic and field observations
associated with sequence development and the position of subaerial exposure surfaces within the
Reelsville unit. Many of the excusion patterns for the parameters tested are more distinct for
particular horizons than others, which is a reflection of several factors including subaerial
exposure duration and degree of diagenetic alteration of the sediment, differential colonization of photosynthesizers on subaerial exposure surfaces, climate, the hydrologic properties of the
sediment at the time of exposure and the degree of overprint on underlying sediment by later
diagenetic events. Based upon the data, six horizons were identified in the Reelsville unit that
corresponds to probable locations of subaerial exposure. Table 4.3 is a summary of geochemical,
physical and statistical evidence for subaerial exposure for each of these six horizons. 127
Figure 4.14: Statistical analysis of δ18O values. The green box is the upper group and the orange box the lower group of values analyzed for each horizon.
128
Discussion
Field observations and high resolution facies analysis of petrographic thin sections suggests the presence of three parasequences characterized by shallowing upward cycles of inner ramp facies within the Reelsville unit and C9 sequence. The base of the first parasequence (P1) begins immediately above the rhizolith dolomitized mudstone. Diverse faunal elements such as brachiopod and echinoderm fragments suggest the base of the first parasequence was deposited under normal marine salinity conditions. The presence of well-sorted peloids, forams and thin micritic coatings on the surface of grains and the underside of brachiopod fragments suggest deposition within a transitional shoal-lagoon environment. The overlying ostracod wackestone is consistent with a restricted lagoon setting. The top of the first parasequence is capped by a massive dolomitized mudstone deposited in a tidal flat environment. The dolomite consists of very fine-grained rhombs suggesting penecontemporaneous formation (Scholle and Ulmer-
Scholle, 2003). The base of the second parasequence (P2), which occurs at 48.3 meters, is a skeletal wackestone dominated byostracod, echinoderm and bachiopods. fines upward into a dolomite-rich mudstone. The mudstone contains very fine-grained dolomite rhombs, fenestral pores and minor quartz grains suggesting shallow water restricted marine deposition. The second parasequence is interpreted as being deposite in a restricted lagoon-tidal flat setting. The base of the third parasequence (P3) is a spiculiferous-ostracod wackstone. The amalgamation of pellets creates a clotted fabric in the mud matrix. Fenestral pores, fine dolomite rhombs and minor quartz grains are also present. Overlying this unit is a poorly sorted peloidal packstone containing mud intraclasts, fragmented echinoderm and brachiopod grains. Capping the third parasequence is a rhizolith dolomitized mudstone displaying clotted fabric and containing minor 129
Table 4.3: Summary of the geochemical, physical and statistical evidence for the presence of subaerial exposure (Horizons 1-6) found within the Reelsville unit at Jellico, Tennessee.
129 130 sub angular quartz grains which transitions into a 10 cm thick collapse breccia horizon. The C8-
C9 sequence boundary is coincident with the top of the P3 boundary. The P3 was deposited in a restricted lagoon setting.
The C8-C9 sequence boundary is characterized by a high dolomite concentration below the boundary and a low concentration above. Calcite is the dominant mineral in facies interpreted as normal marine, whereas facies containing low faunal diversity, suggesting restricted marine conditions, contain high dolomite concentrations. Dolomite concentrations are greatest towards of the top of parasequence boundaries where facies interpreted as tidal flat prevail. The precipitation of fine-grained dolomite in the pore spaces of sediment directly from hypersaline marine water is common in tidal flat environments of the Arabian Gulf region, the modern analog for the late Meramecian-middle Chesterian (Chafetz and Rush, 1994; Illing and Taylor,
1993; Illing and Wells, 1964; Mackenzie, 1981; Shinn, 1983; Warren, 2000; Warren, 1991;
Wenk et al., 1993). Overall, there is an increase in dolomite concentration from the base of the
Reelsville unit towards the C8-C9 sequence boundary suggesting that marine salinity increased during deposition of the upper C8 squence. The spike in dolomite concentration at the base of the Reelsville unit (46.3 m) coincides with the restricted marine facies associated with the surface of subaerial exposure. Spikes in dolomite concentration at 46.8 and 48.4 m, which coincide with physical and geochemical signatures suggesting subaerial exposure, may be related to the hypersaline reflux through the sediment immediately below during periods of exposure.
The absence of distinct spikes and the overall dolomite concentration high between 49.3 and 51.1 m is most likely the result of diagenetic overprinting resulting from the greater degree of hypersaline reflux through the upper Reelsville during subaerial exposure within the interval and the greater duration of exposure associated with the development of the C8/C9 boundary. 131
Quartz concentration remains constant at 4 to 6% through the lower half of the Reelsville unit and then increases to 6 to 19 % above 49 m. The quartz concentration high in the 2.4 meter region below the C8-C9 boundary is possibly the result of the combined effects of a greater windblown detrital input and differential dissolution of carbonate at surfaces of subaerial exposure. Quartz concentrations above the C8/C9 boundary are between 6 and 9 %. Data not presented in this study show that the percentage of quartz remains fairly constant through the lower half of the C9 sequence but attain values between 15 and 29 % at the C9/C10 boundary.
The relatively higher concentration of quartz in the lower C9 sequence compared to the upper C8 and lower half of the Reelsville may represent a greater siliciclastic input resulting from increased weathering of exposed rocks surrounding the basin during the onset of more humid conditions. This conclusion is supported by results from Al-tawil and Read (2003) who, based upon lithofacies analysis, suggest that the C8/C9 boundary marks a transition from semiarid-arid conditions to more humid conditions during the Chesterian epoch.
One very noticeable trend in the clay mineralogy data is the presence of corrensite within the Reelsville unit and its absence in the underlying strata (upper Beaver Bend Formation) and immediately above the C8/C9 sequence boundary (lower Beech Creek Formation). Precursor minerals to corrensite formed in hypersaline tidal flat and restricted lagoon settings of the upper
C8 sequence and later underwent neoformation to form corrensite during mesodiagenesis. This in conjunction with relatively high dolomite concentration and positive excursion patterns in mean δ18O value, suggests that the Reelsville unit was deposited in a highly evaporative semiarid-arid environment. The presence of corrensite in strata that represent normal marine
132 conditions that underlies tidal flat deposits suggests that hypersaline reflux extended into these sediments which resulted in the diagenesis of pre-existing clay minerals into Mg-rich precursor minerals to corrensite.
The Reelsville unit is characterized by a decrease in the mean δ13C value below and an
increase in mean δ18O value at or below surfaces of subaerial exposure. The negative excursion
13 in mean δ C value is consistent with the incorporation of carbon from soil gas CO2 that is
relatively enriched in 12C in diagenetic. This pattern parallels that observed in Pleistocene limestone in Barbados (Allan, 1976; Allan et al., 1977; Allen and Matthews, 1982), Jurrassic limestones in Louisiana (Humphrey et al., 1986), upper Pennsylvanian-Permian limestones in
West Texas (Saller et al., 1994; Saller et al., 1999; Saller and Henderson, 1998), Mississippian
limestones in Virginia (Allen and Matthews, 1982) and Ordovician limestones in Tennessee
(Railsback et al., 2003; Theiling et al., 2007). The concomitant positive increase in mean δ18O
value is consistent with evaporative enrichment of 18O in porewaters at the subaerial exposure
surface where water containing 16O goes into the gas state preferentially over water containing
18O. This pattern has been observed in δ18O values for carbonates in various settings (Allen and
Matthews, 1982; Railsback et al., 2003; Theiling et al., 2007). Evaporative enrichment of 18O in pore water at surfaces of exposure during the deposition of the C8 sequence is consistent with paleoenvironmental interpretations suggesting arid-semiarid climate during the middle
Chesterian (Al-Tawil, 1997; Al-Tawil and Read, 2003; Al-Tawil et al., 2003; Smith et al., 2000;
Smith and Read, 2001; Smith and Read, 1999; Wynn and Read, 2006).
An interesting aspect of the data is the magnitude of the excursions in mean for both δ13C when compared with results observed by Theiling et al. (2007) in Ordovician limestones of
Tennessee. The two greatest negative excursions in mean δ13C within the study interval occur at 133
45.7 and 50.8 m which represent a 1 and 3 ‰ difference respectively with the underlying Beaver
Bend and overlying Beech Creek Formations. The maximum difference in mean δ13C at 50.8 m is 0.5 ‰ greater than the maximum excursion observed by Theiling et al. (2007) below the
M4/M5 sequence boundary. This difference in mean δ13C value between the two studies can be
contributed to the extent and type of plant cover that existed on subaerial exposure surfaces
during the Ordovician compared to the Mississippian. The abundance of vascular plants that
prevailed during the Mississippian compared to the Ordovician would have favored the
12 enrichment of C in soil gas CO2 and ultimately into diagenetic carbonates below surfaces of
subaerial exposure to a much greater extent. Root systems of vascular plants have the capacity
12 to penetrate the sediment surface more deeply facilitating the enrichment of C in soil gas CO2
below the sediment surface
The magnitude of the excursion in mean δ18O in the study interval is much greater than
that seen by Theiling et al. (2007). The two greatest positive shifts in mean δ18O occur at 46.3 and 50.2 m which represent a 4 and 2.5 ‰ difference respectively with the underlying Beaver
Bend and overlying Beech Creek Formations. In contrast, the maximum positive excursion in mean δ18O observed by Theiling et al. (Theiling et al., 2007) was approximately 1 ‰. This large
difference in excursion between the two studies can be attributed to climate. The middle
Chesterian of the Mississippian in the Appalachian Basin is characterized as semiarid-arid (Al-
Tawil and Read, 2003) whereas the Ordovician carbonates of Tennessee were deposited under
relatively more temperate conditions. As such, the evaporative fractionation potential for 16O bearing water would have been greater during the Mississippian compared with the Ordovician.
Maxima in variance of δ13C and δ18O values of samples correspond with subaerial
exposure surfaces at the base of the Reelsville unit, the tops of shallowing upward parasequences 134 and below the C8-C9 sequence boundary. Theiling et al (2007) attribute maxima in variance for
δ13C and δ18O samples they observed in Ordovician limestones to local patchiness of vegetation,
variation in topography of the sediment surface, lateral variability in porosity and permeability
caused by cracks, burrows and borings in sediment as well as lateral variation in the original
mineralogy and porosity and permeability of the sediment. In addition, zones of high variance at
flooding and transgressive surfaces were attributed to the mixing of diagenetically unmodified
post-exposure sediment with modified pre-exposure sediment (Theiling et al., 2007). Increases
in variance at surfaces of exposure observed in this study are consistent with patterns observed in
Ordovician limestones of Tennessee studied by Theiling et al. (2007). They found that replicate
sampling from three horizons (ten separate samples collected 2 m apart per horizon) below the
M4/M5 sequence boundary (Holland and Patzkowsky, 1998) had ranges of δ13C value between
2.4 (variance of 0.6) and 1.2‰ (variance of 0.10) and ranges of δ18O value between 2.8 ‰
(variance of 0.8) and 1.4 ‰ (variance of 0.15). In contrast, replicate sampling (ten samples taken
from a single hand sample) from excursion horizons within the Reelsville unit had ranges of δ13C value between 1.5 (variance of 0.15) and 0.8 ‰ (variance of 0.05) and ranges of δ18O value
between 3.6 (variance of 1.12) and 1.1 ‰ (variance of 0.16). High variance with respect to δ13C in the Ordovician samples is probably the result of plant colonization patchiness on the subaerial exposure surface. In contrast, the relatively low variance with respect to δ13C in the Reelsville
may have been a function sampling. Because ten subsamples were collected from an individual
hand sample the effects of lateral patchiness would have been minimized resulting in lower
variance values compare to those seen by Theiling et al. (2007).
There is one region within the Reelsville unit that does not follow the variance pattern
seen in the Ordovician carbonates of Tennessee. At a position of 51.2 m, the largest variance in 135
δ13C value does not appear to correspond to maxima in variance in δ18O of samples despite the
large positive excursion in mean δ18O and δ13C. In accordance with results of Theiling et al.
(2007), there should be a corresponding high in δ18O variance; however, none exists. Field
observations show that within this portion of the stratigraphic section is a region of diagenetic
alteration based upon the following characteristics: 1) the rock is white-tan in color, brittle and
appears leached; 2) it has a massive texture and is chalky and; 3) Fenestrae are visible to the eye
and in petrographic examination. These data suggest that diagenetic alteration of pre-existing
carbonate was sufficiently thorough with respect to δ18O that variance in samples within this
horizon is very low.
Sample variance for δ13C and δ18O is not related to the total surface area of the slabbed
surface from which the samples were collected. The slabbed surface of each sample was gridded into ten areas that were approximately 2 cm2. One sample from each 2 cm2 grid area was taken
from the surface. However, some hand samples were larger than others resulting in the total
surface area from one hand sample to another varying slightly. To assess if the variance
observed in the data might be related to the small difference in total area between hand samples,
binary plots of variance in δ13C and δ18O values of samples vs total sample area were generated
to see if any relationship existed between these variables. The corresponding binary plots
demonstrate a lack of correlation for these two variables (Figure 4.15). Correlation coefficients
calculated between variance in δ13C and δ18O of samples and total sample area returned values of
-0.30 and -0.17, respectively. These data suggest that the scale of the surface area sampled did not affect variance, and that sample size was sufficiently large to account for the geochemical heterogeneity present below surfaces of subaerial exposure.
136
Figure 4.15: Binary plots of variance vs sample area from samples of the upper C8 and lower C9 sequences from Pine Mountain locality: A) δ13C variance vs total sample area; B) δ18O variance vs total sample area.
137
Other Surfaces of Subaerial Exposure
The most recent sequence stratigraphic interpretation of the Reelsville unit (Al-Tawil and
Read, 2003) identified two surfaces of subarial exposure, one surface at the base of the Reelsville unit and the other at the C8/C9 boundary. Results from whole rock mineralogy, stable carbon and oxygen isotopes and petrographic analysis combined with field observations in this study suggest that there were six surfaces of subaerial exposure within the Reelsville unit. Criteria for identifying and confirming surfaces of subaerial exposure for each horizon are summarized in the following:
In the field, Horizon 1 is characterized by an obvious truncation feature in the strata.
Student’s t-test for both δ13C and δ18O values confirm that the means of the two horizons are
statistically different. F tests for variance in δ13C and δ18O values also confirmed that there is a
statistical difference between the horizons tested. Petrographic evidence includes root
petrifaction suggesting extensive plant colonization of the subaerial surface. XRD did not
confirm the presence of the clay mineral corrensite. The absence of corrensite in this horizon
could have been the result of one the following: 1) Sulfate reduction was high enough for dolomite precipitation to proceed which would have kept the Mg/Ca ratio low preventing the formation of corrensite precursor minerals; 2) pedogenic processes altered corrensite precursor minerals to nonprecursor forms or; 3) hypersaline conditions never reached levels that would have initiated the formation of corrensite precursors. Analysis of whole rock mineralogy showed that dolomite spiked to 89% below this surface suggesting that conditions of hypersalinity did reach levels sufficient for dolomite precipitation.
In the field, Horizon 2 is characterized by wavy contacts separated by shale partings suggesting subaerial exposure. Student’s t-test did confirm a statistical difference in mean for 138
δ13C values and a statistical difference in mean for δ18O values. F tests for variance did not
confirm a statistical difference for δ13C; however, it did confirm δ18O values tested are different.
Petrographic evidence for a subaerial exposure surface was not observed. XRD did not confirm
the presence of the clay mineral corrensite suggesting that conditions during deposition or
subsequent diagenesis were not conducive to corrensite precursor formation. Whole rock
mineralogy showed that dolomite spiked to 19% below suggesting that evaporative conditions
existed to an extent that dolomite precipitation was able to proceed. The proximity of Horizon 2
in relation to Horizon 1 might suggest that the geochemical excursion patterns in Horizon 2
might be the result of the incorporation of diagenetically altered material from Horizon 1 mixing
with unaltered sediment; however, there was no evidence of an erosional lag deposit above
Horizon 1 during field observation or in the sample collected. Furthermore, petrographic
evidence does not show the presence of any intraclasts.
Field observations record two erosional surfaces separated 30 centimeters apart at the
Surface 3 location. Student’s t-test did not confirm a statistical difference in mean for δ13C values; however, a t-test did confirm a statistical difference in mean for δ18O values. F tests for
variance in δ13C and δ18O values confirmed that the horizons tested are statistically different.
Petrographic evidence could not confirm the presence of a subaerial exposure surface. XRD did
confirm the presence of the clay mineral corrensite and whole rock mineralogy showed that
dolomite concentration spiked to 44%.
Field observations did not reveal the presence of an erosional surface at the location of
Horizon 4 suggesting the surface is cryptic in physical expression. Petrographic analysis provides evidence for deposition in tidal flat environment which may have experienced subaerial exposure. Student’s t-test for both δ13C and δ18O values confirmed that there is a high probability 139 that the means of the two horizons are statistically different. F tests for variance for δ18O values
tested confirmed that there is a statistical difference but this was not true for δ13C. XRD results
confirm the presence of the clay mineral corrensite and whole rock mineralogy showed that
dolomite concentration increased to 80%, which suggest highly evaporative marine conditions.
Field observations record an erosional surface at Horizon 5. Student’s t-test for both δ13C and δ18O values confirmed that there is a high probability that the means of the two horizons are different. The F test for variance in δ18O values did not confirm a statistical difference between
the two horizons. The low variance relative to δ18O values suggests that diagenetic alteration of
sediment by meteoric water may have been substantial below this surface. Petrographic evidence
suggests a tidal flat depositional environment and fenestral pores with geopetal structure suggest
periodic wet-dry cycles. Large excursions in mean δ13C and δ18O values and high variance in
δ18O values suggests this surface experienced high a high degree of diagenetic alteration.
However, if diagenetic alteration of carbonate was thorough below this surface, one might expect
the variance in δ13C values to be low as well. High variance in δ13C values probably resulted
from restricted plant colonization at the surface due to arid conditions, which is suggested by the
large positive excursion in the mean of δ18O values, resulting in limited 12C input into the
diagenetic system and less than complete incorporation into diagenetic carbonate. In addition, petrography did not reveal the presence of any root petrification, which suggests plant colonization was limited to non-existent. The f-test for variance in δ13C values did suggest that
there is a high probability that the horizons are different. XRD results confirm the presence of
the clay mineral corrensite and whole rock mineralogy showed that dolomite concentration
maintained about 82% suggesting arid conditions continued. 140
Horizon 6 is coincident with the C8/C9 sequence boundary. Petrographic observations identified spar-filled elongate and circular tubular voids indicative of root petrifaction. Student’s t-test for both δ13C and δ18O values confirmed that there is a high probability that the means of
the two horizons are different. F-tests for variance in δ13C and δ18O values also confirmed that
there is a high probability that the horizons tested are different. Fine-grained dolomite crystals in
mud matrix and minor ostracod elements suggest the depositional environment was a paleosol
developed on top of lagoon or interswale deposit. XRD results confirm the presence of the clay
mineral corrensite and whole rock mineralogy showed that dolomite concentration dropped from
a high of 83 to 4% across the C8/C9 boundary suggesting that a rapid transition from highly evaporative to normal marine conditions.
Conclusions
High resolution petrographic, mineralogic and geochemical analysis of limestone can
improve our understanding of sediment diagenesis within a sequence stratigraphic framework.
Clay mineral and bulk rock mineralogy can be used to constrain paleoenvironmental conditions
during deposition. Increasing dolomite concentration within the Reelsville unit suggests that
marine conditions became more restricted towards the C8/C9 boundary. The presence of
corrensite within the Reelsville is further evidence of restricted marine depositional conditions.
Spikes in dolomite concentration correlate with other physical and geochemical evidence
indicative of subaerial exposure suggesting that it might prove to be a useful proxy for
identifying such surfaces.
Results from this study suggest that geochemical variability in the physical expression of
δ13C and δ18O values in carbonate below surfaces of subaerial exposure may be sufficiently large
in relatively small samples. Geochemical heterogeneity in δ13C and δ18O values at the scale of a 141 hand sample reveals the presence of exposure surfaces identified at broader spatial scales. It may also reveal cryptic exposure surfaces that cannot be recognized in the field. Further work needs to be undertaken to determine if those cryptic surfaces are recognizable at larger spatial scales.
Furthermore, methods described in this study could easily be applied by researchers whose only sample source are cores. The relatively small sample size, 100-300 micrograms of sample in the case of δ13C and δ18O analyses and 10 to 15 grams for petrography and clay mineral analysis,
suggest that a robust analysis from a comprehensive data set can be obtained with minimal
sample size.
142
CHAPTER 5
SUMMARY OF RESULTS
The major objectives of this study outlined in Chapter One were as follows: 1)
Characterize geochemically the Mg-smectite to Mg-chlorite transition and describe the factors that affect the process; 2) identify diagenetic minerals that convey relevant information concerning the position of sequence boundaries; 3) determine whether heterogeneity in geochemical expression of stable carbon and oxygen isotopes exists at the scale of a hand-sized sample and trace the temporal variations in mineral and stable carbon and oxygen isotope composition to assess climatic conditions during deposition. To this end, the following conclusions of this study are as follows; 1) Analytical data from 163 microprobe analyses from ten different studies indicate that the Mg-smectite to Mg-chlorite transition is characterized chemically by a decrease in alkali and alkaline earth metals with decreasing Si/(Si+Al) and an increase in Al+3 for Si+4 substitution in tetrahedral sites (Figure 7, Chapter Two). Ternary plot of
%Si, %Al and %Mg+%Fe+%Mn in mafic phyllosilicates clearly show a decrease in %Si and increase in %Al and %Mg+%Fe+%Mn during the smectite to chlorite through corrensite transformation (Figure 5, Chapter Two). This increasing trend in substitution suggests that the
availability of Al+3 may be an important factor in how the transition proceeds and which phase
predominates (Hillier, 1993). Increases in Fe, Mg and Al during the transition are attributed to
the increase in hydroxide fixation during prograde neoformation. These trends in geochemistry
are consistent for both igneous and sedimentary diagenetic environments. Variability in the
Fe/(Fe+Al) of octahedral cations appears to be dominated by whole-rock composition of 143 hydrothermally altered mafic igneous rocks; however, this phenomena does not appear to hold true for Si/(Si+Al) under similar diagenetic environment (Bettison-Varga and MacKinnon,
1997).
Several factors contribute to the smectite to chlorite transformation through corrensite.
According to data summarized from the literature, a range of stability for corrensite exists between 60 and 300oC for a variety of lithologic and diagenetic settings. The transition appears
to be affected by substrate diffusive capacity and resulting fluid/rock ratios where discontinuous
progression of smectite to chlorite through corrensite occurs during a high diffusive regime
whereas continuous progression, a gradual change involving randomly and regularly interlayered
metastable structures, is associated with diagenetic microenvironments exhibiting low diffusive
capacity and resulting mineral disequilibrium conditions. To what degree diffusive capacity
plays in the smectite to chlorite transformation in sedimentary rocks remains undocumented.
Organic-inorganic interactions promoting reducing conditions appears to promote the formation
of the chlorite phase, which results from increased solubility and availability of reduced iron to
occupy octahedral sites; however, the literature does not report to what extent this factor plays in
the smectite to corrensite transformation process. Inorganic-Inorganic interactions (e.g. reactions
between phyllosilicates and Mg-bearing minerals such as dolomite) might be a factor in the
occurrence of corrensite, however, data reported in the literature is not in full support of the
antithetic relationship between corrensite and dolomite described by Hillier (1993). Indeed,
results from this study suggest that corrensite can exist in rock with high concentrations as well
as low concentrations of dolomite. The lack of consensus in the literature and conflicting results
from this study suggest that inorganic reactions during burial diagenesis may not be sufficient to 144 explain the results observed by Hillier (1993); competitive inhibition of dolomite over corrensite precursor minerals or visa versa early on in the diagenetic process may prove to be a more significant factor in the resulting mineral-lithofacies relationship (Jeans et al., 2005).
One cautionary aspect of the research identified during data analysis was the lack of information concerning the percentage of expandable layers reported in the literature. Applying the formula for the percentage of expandable layers in described by Hillier (1993) is necessary for characterizing chlorite-smectite minerals. For example, many analyses reported as
“smectite”, “corrensite” and “chlorite” are not correct when the percentage of expandable layers is calculated. Constraining the definition of chlorite-smectite minerals to 0-10% expandable layers for chlorite, 45-55% expandable layers for corrensite and 90-100% expandable layers for smectite was important in defining geochemical groupings for each respective mineral. Binary
plots of interlayer cation totals (Na+Ca+K) against Si/(Si+Al), tetrahedral Al total against Si/(Si+Al)
and a ternary plot of %Si, %Al and %Mg+%Fe+%Mn in mafic phyllosilicates illustrate how
geochemical groupings for smectite, corrensite and chlorite cluster together when mineral phase is
constrained to % expandable definitions (Figures 5, 6 and 7, Chapter Two). In the absence of this
constraint, geochemical trends appear as a continuum between smectite to chlorite lacking
discrete groupings. In absence of this additional information, the smectite to chlorite
transformation through corrensite may appear as a continuous reaction series instead of
discontinuous (the dominant transformation mechanism observed in HRTEM studies). In
addition, corrensite would not appear as a discrete phase in the continuum.
2) Corrensite proved to be a successful proxy for sequence boundary identification in
Mississippian aged carbonates deposited under semiarid-arid conditions for sequences that were
deposited under sufficient accommodation to allow for normal marine conditions to prevail.
XRD analysis of clay minerals separated from limestones from a sequence in the Tuscumbia- 145
Monteagle Formations in northwest Georgia reveal the presence of corrensite in association with restricted marine deposits in the vicinity of sequence boundaries. Normal marine deposits of the transgressive systems tract and early high stand systems tract contained smectite, illite and minor kaolinite. Corrensite appeared deeper within each succeeding parasequence of the late high stand systems tract deposits suggesting that marine waters became more saline during these shallowing events prior to sequence boundary formation. This pattern was particularly evident in the two parasequences (P7 and P8 of Figure 10, Chapter Two) immediately below the inferred position of sequence boundary; corrensite appears in P7 at the very top of the shallowing upward cycle and the upper half of the succeeding P8 parasequence.
In contrast to the normal marine sequence, corrensite was the dominant clay mineral phase throughout the restricted marine sequence of the Saint Louis Formation in Jellico,
Tennessee. This sequence was less than half the thickness of the sequence of the Tuscumbia-
Monteagle Formations suggesting deposition under less accommodation space. In addition, this sequence was dominated by restricted marine and tidal flat facies. Corrensite appears both in the transgressive systems tract and high stand systems tract of this sequence. However, corrensite quickly disappears at the transgressive surface above the M1/C1 boundary (Figure 11, Chapter
Two). There is one sample above the M1/C1 boundary that contained poorly ordered corrensite but that could be attributed to erosion of the upper M1 sequence during relative sea-level rise.
These data illustrate that a significant difference in the physical expression of the mineral corrensite exists in sequences deposited under different accommodation regimes, which may prove useful in sequence delineation. For example, the strata below the C8/C9 sequence boundary (Reelsville unit) in Jellico, Tennessee are complex in terms of facies relationhips.
Based upon petrographic evidence alone, the data suggests it was deposited under limited 146 accommodation space represented by inner ramp lagoon and peritidal mudflat environments.
Geochemical, mineralogic and petrographic evidence clearly indicate that the base and top of the
Reelsville unit is bounded by subaerial exposure surfaces. Negative excursion in the average stable C isotope value is indicative of the colonization and decay of photosynthesizers at these bounding surfaces. Positive excusion in stable O isotope value is indicative of evaporative fractionation and concentration of 18O in diagenetic carbonate at surfaces of exposure. In
addition, dolomite concentration highs associated with bounding surfaces suggest evaporative
concentration of shallow marine waters. Petrographic evidence indentifies the presence of root
petrification in the lower bounding surface and brecciation in the upper surface. In addition,
corrensite is a ubiquitous clay mineral phase throughout the Reelsville unit similar to the M1
sequence. In terms of thickness, the Reelsville unit is approximately 5 meters thick, which is
very similar to the M1 sequence (4 meters thick). Based upon this evidence and similarities to
the M1 sequence, the Reelsville unit could represent a stand alone sequence deposited under
limited accommodation space. This would suggest that the C8 sequence of Al-tawil and Read
(2003) may actually represent two separate sequences.
3) Previous workers attempting to identify surfaces of subaerial exposure with excursion
patterns in δ13C and δ18O values assumed that interstratal lateral geochemical heterogeneity was
small enough to be ignored. Theiling et al. (2007) showed that geochemical heterogeneity exists
between replicate samples taken along a single stratigraphic horizon below surfaces of subaerial
exposure in Ordovician limestones of Tennessee. These results highlight potential errors made by previous workers attempting to identify surfaces of subaerial exposure with single samples collected along a single vertical transect. Theiling et al. (2007) recognized that geochemical
heterogeneity can exist in replicate samples collected from a single stratigraphic horizon 147 separated two meters apart, however, the scale of that heterogeneity was never explored. Could the geochemical heterogeneity observed in samples spaced meters apart along a single stratigraphic horizon be observed in replicate samples taken from a single hand-sized sample?
In terms of facies relationships, the Reelsville unit at Jellico, Tennessee is a complex disconformity bounded unit that was deposited under limited accommodation space and represented by inner ramp lagoon and peritidal mudflat environments. The well developed sequence boundary capping the Reelsville (the C8/C9 boundary) and complex nature of deposition is believed to be the result of a significant sea-level fall associated with a third-order super sequence cycle (Al-Tawil and Read, 2003). Field observations coupled with petrographic analysis reveal three shallowing upward parasequence cycles. However, seven physical features were identified within the parasequence cycles as potential regions of subaerial exposure. Data from replicate sampling of single hand-sized samples collected at 30 cm intervals along a single vertical transect exhibits sufficient geochemical heterogeneity in δ13C and δ18O values to confirm
the position of six surfaces of subaerial exposure in the Reelsville unit, which correspond the physical features suggestive of subaerial exposure. Student’s t-test and f-test results of δ13C and
δ18O values confirm statistically the presence of six surfaces of subaerial exposure in the
Reelsville unit where only two had previously been identified. Geochemical heterogeneity in
δ13C and δ18O values in hand-sized samples imply that replicate sampling spaced meters apart
along stratigraphic horizons may be unnecessary to attain statistical verification of surfaces of
subaerial exposure. If geochemical variation in δ13C and δ18O is large enough to be recognized
in relatively small samples then replicate sampling technique describe in this paper could
possibly be applied by researchers who are limited to sampling material from cores .
148
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APPENDIX A
Microprobe analyses of expandable mafic phyllosilicate minerals
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APPENDIX B
X-ray diffraction of bulk powder
M1-C1 Sequence
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Sample ID: M1-11 Type: Bulk Rx Powder Location: Jellico, TN Strat Position: 2.90m
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XRD of Bulk Powder Continued
Lookout Mountain Sequence
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240
241
242
243
244
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248
249
250
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252
253
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XRD of Bulk Powder Continued
Upper C8 and Lower C9 Sequence
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APPENDIX C
X-ray diffraction of clay mineral (<2µm) fraction (Red=Ethylene Glycol, Black =Dry)
M1-C1 Sequence
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301
302
303
304
305
306
307
308
309
310
311
312
313
XRD of Clay Mineral (<2µm) Fraction Continued
Lookout Mountain Sequence
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318
319
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321
322
323
324
325
326
327
328
329
330
331
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333
334
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336
337
338
339
340
341
342
343
344
345
346
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348
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352
353
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355
XRD of Clay Mineral (<2µm) Fraction Continued
Upper C8 and Lower C9 Sequence
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361
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APPENDIX D
Summary of δ13C results from the upper C8/C9 sequences
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Summary δ18O results from the upper C8/C9 sequences, Jellico, Tennessee.