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dissertation entitled

Dolomitization and Porosity evolution

presented by

Michael A. Miller

has been accepted towards fulfillment
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m‘ ----

DOLOHITIZATION AND POFIOSITV
EVOLUTION

5!!

Michael A. Miller

A DISSERTATION I

Submitted to
Michigan State University
in partial fulfillment of the requirements
for the degree of

DDCTDFI or PHILOSOPHY
Department of Geology

I 988

ABSTRACT
DOLOI‘IITIZATION AND POFIOSITY EVOLUTION
By
Michael A. Miller

Dolomite rock texture, and therefore its porosity, is controlled by
the interplay of factors that influence reaction kinetics. These factors
include: 1) the surface area of the substrate, 2) the saturation state of
the solution, 3) temperature, 4) substrate mineralogy, and 5) the role of
inhibiting ions. Dolomite porosity is also affected by the flux of
bicarbonate through the system in that this influences the amount of
dolomite cement that can be precipitated.

In the Trenton dolomites (Michigan Basin) and the Sero Domi
dolomites (Aruba, NA), the surface area of the substrate has a control on
the resultant porosity of the dolomite. The dolomitization of a finely
crystalline (high surface area) limestone, resulted in dense dolomite
nucleation and a dolomite with porosities similar to, or lower than, the
host. Dolomitization of a more coarsely crystalline (low surface area)
limestone, resulted in less dense dolomite nucleation and a dolomite with
porosities higher than the host.

In the Trenton Formation dolomites, the flow direction for the
dolomitizing fluid can be inferred from the distribution of the dolomite
and trace element gradients. The change in fluid chemistry of the

dolomitizing solution across the study area had no significant influence on

the resultant porosity of the dolomites.

Limestones and dolomites with similar textures from the Sero Domi
Formation and the Eniwetok Atoll were studied to assess the role of
bicarbonate flux on the porosity evolution of the dolomites. It was
determined that the f lux of bicarbonate through the sediment was not a

limiting factor in the porosity evolution of these dolomites.

AKNOWLEDGEI‘IENTS

This research was partially funded by both the National Science
Foundation (EAR B2 14106) and the Department of Geological Sciences at
Michigan State University. I wish to thank Eugene Shinn and the USGS at
Fisher Island Station for the use of the drilling equipment on Aruba, NA.
The government of Aruba and the Aruba Mining Company were also most
helpful with equipment and support. A special thanks to Total Petroleum
who granted the use of the Trenton cores and supplied the porosity and
permeability data. Sampling of the Eniwetok core was granted by the
Smithsonian Institution, with sampling directed by Warren Blow. Phil
Koro of Schlumberer logging company was also helpful with the
interpretation of the fracture identification log.

A special thanks to my thesis advisor Duncan Sibley for his support. I
would also like to thank the other members of my committee Dave Long,
Mike Velbel and Tom Vogel.

Thanks to all my fellow students who made my tenure at MSU a
memorable experience. I would especially like to thank, Tim Flood PhD,
Greg Giles, Nadine Romero, Jim Mills, Bill Sack, and Jerry Grantham for
their support. Thanks to the various geochemistry types who helped so
much with the collection of the trace element data and an occasional rod
hockey game especially, Tim Wilson, Dale Rezabeck and Nancy Fegan.

I would also like to gratefully acknowledge the help and support of
Robert Hayes of the MDNR.

As always, I would like to thank my parents Robert and Dorothy Miller
for their unmitigated moral and financial support.
Finally, I would like to thank my wife and best friend Beth for her

unerring support through the past five years.

TABLE OF CONTENTS

ACKNOWLEDGEMENTS . . . . . . . ii
LIST OF TABLES . . . . . . . . vii
LIST OF FIGURES . . . . . . . x
INTRODUCTION . . . . . . . I
PREVIOUS WORK . . . . . . . . 3

MODEL RELATING SUBSTRATE SURFACE AREA
AND SATURATION STATE OF THE SOLUTION

TD DDLDMITE PDRDSITY . . . . . . 5
Model For Dolomite Porosity . . . . 9
Reactant Surface Area in Nature . . . 16
Substrate Surface Area Affects (Examples) . . 19
Ranges of Solution Chemistry in Nature . . 21
Solution Chemistry Affects (Examples) . . 23
METHODS . . . . . . . . . 25

OBSERVED AFFECT S OF SUBSTRATE
SURFACE AREA AND SATURATION STATE OF
THE SOLUTION TO DOLOI‘IITE POROSITY . . . . 27

The Ordovician Trenton Formation
Jackson County, Michigan . . . . 2?

iv

TABLE OF CONTENTS (Cont'd)

Relationship of the Trenton Dolomites
to Fractures.

Discussion of Substrate
Surface Area Affects

Discussion of Solution Chemistry Affects

Pliocene Sero Domi Formation
Aruba, Netherlands Antilles

Discussion of Substrate
Surface Area Affects
BICARBONATE F LUIt

Pliocene Sero Domi Formation
Aruba, Netherlands Antilles

Origin of the Sero Domi Dolomites
Discussion of the Bicarbonate Flux Model
Eocene Carbonates From the Eniwetok Atoll .

Discussion of the Bicarbonate Flux Model

DISCUSSION OF DOLOMITE POROSITY .

SUMMARY .

CONCLUSIONS

APPENDIX I

APPENDIX ll

APPENDIX III

63

?I

78

84

92

94

98

IIO

IIS

II6

I23

I24

126

I28

I30

I36

I48

TABLE OF CONTENTS (Cont'd.)
APPENDIX IV . . . . . . . . I53

REFERENCES . . . . . . . . 158

vi

Table

Table

Table

Table

Table

Table

Table

Table

Table

Table

Table

I0.

II.

LIST OF TABLES

BET specific surface area as a function of grain
size.

The saturation indexes of the solutions with respect
to dolomite

Variation of crystal sizes in Trenton limestone
clasts

Oxygen and carbon isotope analyses of the Trenton
limestones

Minor and trace element compositions (ppm) of
the Trenton limestones.

Dolomite crystal size distribution in the clasts
and the matrix between the clasts from the Trenton.

Oxygen and carbon isotope analyses of the Trenton
dolomites.

Oxygen and carbon isotope analyses of dolomites
from the partially dolomitized cores, Trenton
Formation.

Strontium concentration (ppm) of the Trenton
Formation dolomites.

Manganese, iron and zinc concentrations (ppm) of
the Trenton Formation dolomites.

Stable isotopic and trace element compositions of
some ancient dolomites.
vii

I8

22

38

43

50

54

SS

56

58

6|

Table

Table

Table

Table

Table

Table

Table

Table

Table

Table

Table

Table

Table

I2.

I3.

I4.

IS.

I6.

I7.

I8.

19.

20.

21.

22.

23.

24.

LIST OF TABLES (cont'd)

Comparison of oxygen isotopic signatures between
Trenton dolomite types.

Comparison of manganese concentrations (ppm)
between Trenton dolomite types. .

Comparison of iron concentrations (ppm) between
Trenton dolomite types.

Point counting results for the Trenton dolomite
textural comparison.

Porosity and permeability of the Boo Doei
carbonates (Aruba, NA).

Oxygen and carbon isotope analyses of the Sero
Domi limestones, (Boe Doei).

Stable isotope and trace element analyses of the
Sero Domi dolomites (Boe Doei). .

Darcy velocities and mass fluxes.

Oxygen and carbon isotope analyses of the Sero
Domi limestones (Rooi Hundu).

Stable isotope and trace element analyses of the
Sero Domi dolomites (Rooi l-lundu).

Porosity and permeability of the Rooi Hundu
carbonates (Aruba, N.A.).

Stable isotopic and trace element compositions of
selected Cenozoic dolomites.

Porosities and permeabilities of the Eniwetok
carbonates.

viii

69

70

72

73

87

89

9I

97

I00

IOS

I08

114

I2I

Table AI-l.

Table All-I.

Table AlV-I.

Table AlV-2.

LIST OF TABLES (cont'd.)

Comparison of Compensated Neutron and
Whole Core Porosity of the Trenton Dolomites .

Whole Core Porosity Analyses of the Trenton
Formation Carbonates

Compensated Neutron Log Porosity Comparison
of the Nonconglomeritic Zones in the Trenton
Formation Dolomites

Whole Core Porosity Comparison of the
Conglomeritic Horizon in the Trenton Formation
Dolomites

I3l

I36

I53

I55

Figure I.
Figure 2.
Figure 3.
Figure 4.
Figure 5.
Figure 6.
Figure 7.
Figure 8.
Figure 9.
Figure 10.
Figure ll.
Figure I2.

LIST OF FIGURES
Model of dolomite porosity evolution in a
generic substrate

Model of dolomite porosity evolution in a
generic wackestone .

Schematic cross-section of the Red River
Dolomites, Montana .

Location Map for the Trenton Formation Study
Area in Jackson County, Michigan .

Map of wells that were used in the study.

Study area map with well lease names listed.
Cores used in this study are circled

Generalized isopach map of the Trenton
Formation in the study area.

Top of the Trenton Formation contour map

West to east cross-section A-A' through the
Trenton Formation in the study area.

North to south cross-section B-B' through the
Trenton Formation in the study area.

Rounded quartz grains in the matrix between
dolomitized clasts in the Trenton.

Blocky low magnesium calcite cement filling
molds molds in the Trenton limestone

X

I0

I4

24

28

29

30

31

33

34

36

39

Figure I3.

Figure I4.

Figure 15a.

15!).

Figure 16.

Figure l?.

Figure 18.

Figure 19.

Figure 20.

Figure 2 I.

Figure 22.

Figure 23.

LIST OF FIGURES (COMO)

Cemented fracture in a limestone clast that
is broken off at the edge of the clast.

Composition of the Trenton Formation
limestone clasts.

Trenton dolomite observed in plane light.

Trenton dolomite as viewed in blue-violet

' light.

Comparison of the crystal sizes between
the dolomitized clasts and the dolomitized
matrix between the clasts in the Trenton
Formation.

Large dolomite crystals replacing allochems
and smaller dolomite crystals replacing mud
in the Trenton Formation. .

Late iron-rich dolomite cement in the Trenton
Formation dolomites.

Comparison of porosity between limestone and
dolomite clasts in the Trenton.

Fossil moldic porosity in the Trenton dolomites.

Dolomite clasts with varing porosity in close
proximity.

Rose diagram of the fracture trend in the
Trenton in the study area. .

Nonplanar dolomite crystal replacing calcite
cement in the Trenton

xi

41

47

47

49

52

S3

62

64

65

66

7S

Figure 24.

Figure 25.

Figure 26.

Figure 2?.

Figure 28.

Figure 29a.

29b.

29c.

Figure 30.

Figure 3I.

Figure 32.

Figure 33.

Figure 34.

LIST OF FIGURES (cont'd.)
Oxygen isotopic signature of the dolomites
versus distance from west to east for the
Trenton dolomites

Location map for the Sero Domi Formation outcrops
on Aruba, N.A.

Percent porosity versus depth for the carbonates
from Boe Doei .

Blade shaped cement crystal included in a dolomite
crystal from Boe Doei

Dolomite crystals from the partial and complete
dolomites from Boe Doei.

Typical limestone from Rooi Hundu.

Dolomite from Rooi Hundu showing mimic
replacement of the coralline algae

Foram mold in the dolomites from Rooi Hundu.

Coralline algae fractured around what is
currently a f oram mold from Rooi Hundu.

Percent porosity versus depth for the carbonates
from Rooi Hundu.

Location map for the core used from the
Eniwetok Atoll.

Selective replacement of the muds by dolomite
and the resistance of coarsely crystalline
fossil fragments

Dolomite cementation after calcite
cementation on Eniwetok. .

xii

82

85

86

90

93

101

102

103

104

107

117

118

119

LIST OF FIGURES (cont'd.)

Figure 35. Percent porosity versus depth for the carbonates
from Eniwetok. . . . . . I20

xiii

INTRODUCTION

Dolomitization is a process that is governed by factors that
influence reaction kinetics (Machel and Mount joy, 1986). Nucleation and
growth rates of the dolomite crystals influence the texture and therefore
the porosity of the resultant dolomite. Recent laboratory experiments
have delineated several key parameters affecting reaction kinetics: a)
Mg/Ca ratio in the solution and substrate mineralogy (Gaines, 1980); b)
crystal size of the reactant (Sibley et al, 1987); c) temperature (Katz and

Matthews, 1977),- and d) the effects of inhibiting ions, such as 304 '

(Baker and Kastner, 1981).

In addition to reaction kinetics, the flux of bicarbonate through the
system is an important consideration in the evolution of dolomite
porosity. Bicarbonate flux is important in that it influences the amount of
dolomite cement that can be precipitated.

Most dolomites have very complicated diagenetic histories.
Consequently, In most studies examining dolomite porosity, the factors
of f ectlng porosity are difficult to constrain. This study is significant in
that constraints can be imposed to evaluate the affects of; I) substrate
surface area, 2) saturation state of the solution, and 3) the flux of
bicarbonate on the resultant dolomite porosity. Several hypotheses are
tested in this study; 1) that crystal size of the reactant creates variations
in resultant dolomite porosity, 2) that changes in the fluid chemistry of
the dolomitizing solution create variations in the porosity of the resultant
dolomite and 3) that the mere flux of bicarbonate through the carbonate
system is not the limiting factor in dolomite porosity development.

Samples from the Ordovician Trenton Formation (Michigan Basin), the

2

Pliocene Sero Domi Formation on Aruba (Netherlands Antilles), and the
Miocene carbonates on the Eniwetok Atoll (South Pacific) were used in this
study. Samples from the Trenton Formation were studied because they
contain carbonate conglomerates. The individual clasts in the
conglomerate have a wide range in composition. This provides an
opportunity to examine what influence the variation in substrate texture
has on dolomite porosity within a close proximity. It is highly improbable
that clasts of strikingly different textures immediately adjacent to one
another experienced differences in solution chemistry and temperature.
The Trenton dolomites also provided an opportunity to look at how changes
in solution chemistry might affect dolomite porosity. These changes in
solution chemistry were inferred from the spatial distribution of the
dolomites and trace element gradients.

Samples from the Sero Domi Formation on Aruba, were studied
because they contain dolomites and limestones with similar original
depositional textures. This allows one to determine whether or not
dolomitization creates or destroys porosity. The amount of bicarbonate
transported into or out of the system can then be determined for mass flux
modelling.

The dolomites from the Eniwetok Atoll were studied to see how
fabric-selective dolomitization would affect dolomite porosity. Previous
work on these dolomites also provides constraints on the timing of
dolomitization and the type of dolomitizing solution for mass flux
modelling. Sal ler (1984a) presents a reasonable case for dolomitization
by deep seawater. Through the use of strontium isotopes in the dolomites
he was able to constrain the timing for the dolomitization process. Using
petrographic data obtained in this study and the data provided by Saller

3

( 1984a), one can model the flux of bicarbonate through the original
carbonate substrate.

in the Trenton and Sero Domi dolomites, it will be shown how the
crystal size of the reactant (limestone) produces the textural/porosity
variations observed. Additionally, in the Trenton dolomites, it will be
shown that changes in dolomitizing fluid chemistry did not affect the
porosity of the resultant dolomite. 1n the Eniwetok and Sero Domi
dolomites, it will also be shown that the bicarbonate flux was adequate to

produce the observed porosities.

PREVIOUS WORK

Porosity evolution of dolomites is most commonly explained by
“volume shrinkage', a model attributed to deBeaumont in 1836 (Van Tuyl,
1914). The model was discussed in greater detail by Murray ( 1960) and
Weyl (1960). They called it the 'local source theory', because the source

OT 0032' W88 OSSUMOO to DB derived from "IO COCUZ. being FBDIBCBO. The

model is based on the dolomitization of a limestone via the following

equation:
rig2+ + 2ceco3 ---- cf:2+ + CaI'Ig(CO3)2 (I)
The resultant dolomite will occupy 13 percent less volume than the

original volume of the calcite. This occurs due to the molar volume
difference between two moles of calcite and one mole of dolomite. Since

4

dolomitization via this reaction should result in a porous dolomite, a -
strict application of this model universally does not seem warranted due
to the existence of many low porosity dolomites.

Other models attribute the porosity of dolomites to the dissolution
rate of calcite being greater than the precipitation rate of dolomite during
the replacement process (Murray, 1930 and Landes, 1946). However,
studies of partial and complete dolomitization by Lucia (1962), Jodry
(1969) and Choquette and Steinen (1985), infer that dolomitization
initially decreased porosity and then increased porosity. These studies
are not consistent with the dissolution rate of calcite being more rapid
than the precipitation rate of dolomite.

Wardlaw (1979) suggested that porosity in dolomite can be attributed to
contact inhibition. Contact inhibition is the tendency for crystals to stop
growing when they come into contact. He suggested that contact
inhibition between dolomite crystals is greater than the inhibition
between calcite crystals, which tend to form compromise boundaries and
destroy porosity. This suggestion however, is in direct conflict with the
fact that compromise boundaries and intergrown dolomite crystals are
observed through standard petrographic and SEM examination. Schmoker
and Halley (1982) suggest that dolomites are more resistant to the
'porosity reducing effects of burial" than limestones.

Limestone and dolomite porosities have been compared in many
studies (Lucia, 1962; Jodry, 1969; Schmoker and Halley, 1982; Halley and
Schmoker, 1983; Longman, Fertal and Glennie, 1983; Baum, Harris and
Drez, I985; Schmoker, Krystinik and Halley, 1985; and Choquette and
Steinen, 1985). These studies however, produce no consistent trends in

dolomite porosity data. For example, Longman et al (1983), Baum at al

5

(1985) and Choquette and Steinen (1985) showed that dolomites were
more porous than their associated limestones. Lucia (1962), Jodry (1969)
and Choquette and Steinen (1985), inferred that dolomitization initially
decreased porosity and then increased porosity. Hal lay and Schmoker
(1983) showed that no difference in porosity exists between limestones
and dolomites of the Cenozoic of Southern Florida. Schmoker and Halley
(1982) also show that shallowly buried dolomites from south Florida are
less porous than their associated limestones. Finally, Schmoker et al
(1985) found that limestone reservoirs in the United States are actually
more porous and permeable than dolomite reservoirs.

None of the models presented can account for the wide range of
dolomite porosities observed in nature.

The texture, is. porosity, of non-detrital rocks is the result of
several factors influencing crystal nucleation and crystal growth. Two of
these factors are; 1) the surface area of the substrate, and 2) the
saturation state of the solution. The affects of these two parameters on
dolomite porosity will be addressed in the following section. The
influence of bicarbonate flux on dolomite porosity will be addressed in a

later section.

I'IODEL RELATING SUBSTRATE _
SURFACE AREA AND SATURATION STATE OF THE SOLUTION
TO DOLOI‘IITE POROSITY

If the flux of bicarbonate through the system is not a limiting factor,

the nucleation rate of the dolomite can determine the texture of the

6

dolomite and consequently its porosity. Two factors which can affect
nucleation rate are; I) the surface area of the reactant particles, and 2)
the supersaturation of the solution with respect to the products. How
these two parameters can influence nucleation rate is addressed below.

The heterogeneous nucleation rate of a phase can be related to the
number of active sites (potential nucleation sites) by the following
equation from Christian (1975, sq. 52.34):

I z (N)"3L(kT/h) exp {-0566 + ag’mm (2)

where N = Number of atoms per unit volume which are on dislocation
lines
L = Dislocation density
AB = Critical free energy for homogeneous nucleation
ag’ = Free energy of activation per atom
k = Boltzmann’s constant
h = Plank's constant

T = Temperature

It is assumed that the number of active sites (dislocations) is
directly proportional to the surface area of the reactant, and therefore,
the nucleation rate is directly proportional to the surface area of the
reactant. An increase in the surface area of the reactant will cause a
corresponding increase in the nucleation rate.

The influence of the saturation state of the solution on the
heterogeneous nucleation rate Is observed by substituting the free energy

term for homogeneous nucleation into the preceding equation. The free

7

energy term for homogeneous nucleation is given by the following equation
from Christian (1975, sq. 46.20):

sec = 463n3/27(9'-g')2 (3)

where
a = Surface free energy
n = Shape factor
9 = Gibb's function per atom (I - liquid, v - vapor)

Substituting this equation into equation (2) above one can see that the
nucleation rate is an exponential function of the saturation state.

This relationship can also be seen in the following equation from
Boistelle (1982):

l = Nov exp [-16 8233/3(kt)3(ln3)2 exp ('Bvam’km (4)

where
N0 = no. of solute molesules/ volume of solution

v = frequency with which a critical nucleus becomes
supercritical

B = volume of a molecule inside the nucleus

3 = interfacial surface free energy between the nucleus and the
solution,

S = degree of supersaturation,

Bvdiff = energy barrier encountered by the bulk solution for volume

diffusion from the bulk solution to the nucleus.

8

We lack the necessary data on the surface free energy of the dolomite
nucleus, the shape factor of the dolomite nucleus, and the dislocation
density of the substrate, to apply these equations quantitatively to
dolomitization. Oualitatively however, from equation (2) it can be seen
that a direct relationship exists between the dislocation density and the
nucleation rate. From equations (3) and (4) it can be seen that a linear
relationship also exists between log I and the degree of supersaturation.
The slope of the curve will be determined by factors in the numerator of
the exponent.

In summary, the rate of nucleation is directly proportional to the
number of active sites on the substrate, which in turn is directly
proportional to the surface area of the reactant. An increase in the
surface area of the reactant will therefore, cause a corresponding increase
in the nucleation rate. The rate of nucleation is also an exponential
function of the degree of supersaturation with respect to the product
phase. An increase in the saturation state of the solution will therefore,
cause a corresponding increase in the nucleation rate.

The porosity evolution of a dolomite is also influenced by the
dissolution rate of the parent limestone. From observations in the
partially dolomitized rocks from the Trenton Formation in this study, it
appears that dissolution of the limestone host occurs at the
limestone-dolomite interface and not whole-sale throughout the sediment.
This is evidenced in the partial dolomites in that no partially dissolved
parent mud or allochems are observed. if whole-sale calcite dissolution
was occurring in the host, there should be evidence of dissolution in both
the mud, and the allochems, in the partial dolomites. lf dissolution of the

parent limestone occurs at the limestone-dolomite interface and no

9

dolomite nucleates on the substrate, then dissolution will not occur. This
unreplaced host can then be dissolved out at some later time creating

porosity.

Model for Dolomite Porosity

The model of dolomite porosity proposed has two parts, 1) the
influence of the substrate surface area, and 2) the influence of the
saturation state of the solution. These two situations will be examined in
relation to varying amounts of bicarbonate.

First, if the growth of the nuclei is not limited by other kinetic
factors, than the porosity of the product will be directly related to the
surface area of the reactant. The affect of substrate surface area and
solute flux will first be examined using a generic substrate, and then using
a generic wackestone. In figure I, substrate (a) has a high density of
dolomite nuclei, and substrate (b) has a low density of dolomite nuclei. The
evolution of the porosity to total exclusion of porosity in both substrates
is depicted. Both substrates begin the process with the same porosity. In
an open system with a solution of constant composition and temperature,
and with excess bicarbonate moving through the substrate, the crystals in
both substrates will grow at the same rate. After a short period of growth
(T 1), substrate (a) with the high nucleation density has a lower porosity
than substrate (b) with the low nucleation density. This is due to the
tighter initial packing of the nuclei. After a longer period of growth (T 2),
the porosity of the substrate (a) with the high nucleation density has been

totally occluded, but substrate (b) with the lower nucleation density still

10

Figure 1. Model of dolomite porosity evolution in a generic substrate.
Substrate (a) has a high density of dolomite nuclei and substrate (b) has a
low density of dolomite nuclei. The substrate in both cases has the same
original porosity. As dolomitization ensues, substrate (a) will lose its
porosity more rapidly than substrate (b) due to the greater number of
dolomite crystals growing. At time T2 substrate (a) has lost its porosity,
whereas substrate (b) still retains porosity. With continued

dolomitization, substrate (b) will lose its porosity as well at a later time
(T3).

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has porosity. If substrate (b) remains in the dolomitizing solution, it will
eventually lose its porosity and become occluded as well (T3). The product
in this sample will have a larger crystal size than the product formed from
the nuclei that were more densely packed initially. The density of the
nuclei will then, determine which portions of the substrate will lose their
porosity first, and which portions of the substrate will retain their
porosity the longest. Substrates with a higher number of nuclei per unit
mass of reactant would have more closely packed products and therefore a
lower porosity than the corresponding substrate with less dense nucleation
at any time (T) as the sediment approaches zero porosity.

The dolomitization of finely crystalline reactants (i.e. mud or
cryptocrystalline allochems) will result in dense dolomite nucleation. This
dense nucleation will result in a finely crystalline dolomite with a
relatively low porosity. The dolomitization of a coarsely crystalline
substrate (9.9. microcrystalline allochems) will result in a lower density
of dolomite nuclei. This lower nucleation density is manifested in the
formation of a coarsely crystalline dolomite with a relatively high
porosity.

Considering the same situation with a limited amount of bicarbonate,
initially the substrate with the higher nucleation density will become more
porous due to the more rapid dissolution of the local carbonate to produce
the dolomite. As dolomitization continues however, and dolomite cement is
added, the substrate with the more densely packed dolomite crystals will
result in the lower porosity dolomite.

Second, the porosity of the dolomite is a function of the saturation
state of the solution. The nucleation rate of dolomite is an exponential

function of the degree of supersaturation with respect to the product

13

phase. An increase in the degree of supersaturation will therefore behave
in a similar fashion as an increase in the surface area of the substrate. A
high saturation state will result in a high nucleation rate and therefore a
high density of dolomite nuclei. A solution with a high saturation state
should therefore result in dense nucleation and tighly packed products.
This is similar to the effect produced by a finely crystalline substrate.
Solutions with a high saturation state will cause dense nucleation and will
result in a low porosity product. Conversely, a solution with a low
saturation state will result in a low density of nuclei and a high porosity
product. Therefore at time T (any time before the total occlusion of
porosity) the substrate subjected to the solution with a higher saturation
state will result in a dolomite with a lower porosity than the same
substrate subjected to a solution with a lower saturation state.

The affects of substrate surface area, saturation state of the solution
and time are demonstrated on a generic wackestone in figure 2. In this
wackestone, the carbonate mud portion will have a high density of dolomite
nuclei due to its high surface area. There are also some coarsely
crystalline fossil fragments in this wackestone that will have a lower
density of nucleation sites due to the lower surface area of these grains.
If this wackestone is subjected to a solution that is highly supersaturated
with respect to dolomite, very dense nucleation will occur in the mud
portion of the wackestone due to both the high surface area of the
substrate and the high saturation state. The more coarsely crystalline

portions of the sample will have fewer dolomite nuclei (IIA). Given an

adequate "UN 01 HCUs', the OOIOMItB RUCIOT Will grow ”Didi” BRO Impinge

on one another. The product will be a dolomite with a porosity similar to

or lower than the limestone that it replaced. The dolomite will retain the

14

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

IIIA me m c III 0

,,,,,,

 

 

 

 

 

 

 

 

 

 

 

SATURATION WITH RESPECT TO DOLOMITE DECREASES

Figure 2. Model of dolomite porosity evolution in a generic wackestone.
Where (1) represents the generic wackestone, (II) is the onset of
dolomitization showing the distribution of the nuclei, and (III) is the
completely dolomitized result (diagram from Sibley and Gregg, 1987).

15

original limestone texture (IllA) due to the dense initial nucleation. Due to

the BXIBHIOIIQ OOT’IVBU HCO3', dolomite cementation [“89 OCCUT' ONO DFOOUCB

a dolomite with a lower porosity than the original limestone.

If the same wackestone is subjected to a solution with a slightly

lower saturation state (118) and little extenally derived HCO3', there is

still relatively high nucleation density in the mud portions of the sample,
due to its high surface area, but little or no nucleation in the more coarsely
crystalline portions of the rock. The more coarsely crystalline portions of
the substrate will therefore be much slower to dissolve and remain
unreplaced. The coarsely crystalline substrate will be slower to dissolve
because the dolomite does not nucleate on it. Since dissolution occurs at
the limestone-dolomite contact, if there is no nucleation there will be no
dissolution. These unreplaced fragments can later dissolve, in which case
fossil moldic porosity will result (1118).

If the same wackestone is subjected to a solution with an even lower
saturation state (NC) and very little externally derived bicarbonate, there
is a lower density of dolomite nuclei. This results in large dolomite
rhombs that are not packed as closely as the rhombs in case A (INC), and no
dolomite cementation. A dolomite which has a porosity that is increased
over the precursor limestone is produced. With the low nuclei density in
the host, the fine textural detail of the host is obliterated.

In the first three cases presented, the amount of time was a fixed
quantity. in example D, the affect of a prolonged time in a dolomitizing
solution with a low saturation state is presented. In case D, there are few
nuclei in the substrate due to the low saturation state. This low saturation

state also results in a slow growth rate. If however, this solution is

16

f luxed through the sediment over a sufficient period of time, then the
dolomite rhombs would eventually impinge on one another and result in a
low porosity dolomite (lllD), that shows textural obliteration.

The previous examples were sediments that experienced one
dolomitization event. For each succeeding event of dolomitization a new

set of factors needs to be analyzed.

Reactant Surface Area in Nature

A mudstone will have a higher number of possible nucleation sites per
unit volume than a coarse crystalline grainstone due to the higher surface
area of the micrite per unit volume. Therefore, dolomite should nucleate
more rapidly and more densely in mudstones as compared to grainstones.
This assumes that the number of active sites on a crystal surface is
directly proportional to its surface area. There are no data available on the
surface area of natural reactants in situ. This appears to be a good working
hypothesis however since Bartlett (1984) conducted dolomitization
experiments with substrates of different surface area, and found that
dolomitization did indeed occur more rapidly in substrates with a higher
surface area to volume ratio. This is also consistent with the common
observation that mudstones and the mud in packstones are often
preferentially dolomitized (Chilinger, 1956; Powers, 1962; Murray and
Lucia, 1967; Armstrong, 1970; Wardlaw, 1979; Ruzyla and Friedman, 1985;
Gawthorpe, 1987).

The magnitude of the surface area affect of nonallochems on
nucleation rate can be estimated by comparing crystal size differences of

17

micrlte and microspar. The size range of micrite is 1-4 microns and the
size range of microspar is 4-50 microns (Bathurst, 1975). If one assumes
that the crystals are cubes, the difference in surface area between micri te
and microspar is approximately four orders of magnitude from 6 microns to
15,000 microns. The magnitude of the surface area affect of allochems on
nucleation rate is difficult to estimate. The sizes of allochems and cement
however can vary by much more than size differences observed between
micrlte and microspar. The crystal sizes of allochems and cement can
range from microns to millimeters.

Crystal size differences occur not only between mudstones and
grainstones, but also between different fossil types. Walter and Morse
( 1984) have shown that fossil types have a wide variety of BET surface
areas, and that these surface area measurements can vary by at least an
order of magnitude (T able 1). Differences in the reactivity of these fossils
to dolomitization are contingent upon not only differences in crystal size,
but also to differences in mineralogical composition, chemical composition
and microstructure (Sibley, 1982; Bullen and Sibley, 1984). From
petrographic observation and experimental studies one can see the
differences in behavior of different fossils to dolomitization. For
example, certain fossils are replaced mimetically, nonmimetically, not
replaced at all, or entirely dissolved out (Murray, 1960; Sibley, 1982;
Bullen and Sibley, 1984). If these effects were due entirely to differences
in crystal size, the following would occur. Red algae which have
cryptocrystalline calcite skeletons (i.e. high surface area) are more likely
to be replaced mimetically, whereas some molluscs which have
microcrystalline calcite skeletons (i.e. low surface area) are more likely to

be nonmimetically replaced or remain unreplaced. This difference in

18

TABLE 1. BET SPECIFIC SURFACE AREA AS A FUNCTION OF GRAIN SIZE

    

 

Observed Observed/Geo-
Specific metrically
Median Grain Surface Predicted
Diameter Area Surface

Grain Type (Microns) (M294) Area
Echinoid 81 0.14 5.7
275 0.09 I 2.4
5 1 3 0.08 20.5
Coral 5 1 0.23 5.9
B 1 0.22 8.9
275 0.17 23.4
5 13 0. I 2 30.8
flo/imede B 1 2.04 82.6
2 15 2. 10 225.8
5 13 2. 1 1 541 .2

 

(Walterand Morse,1084)

19

textural response is best explained by the different number of potential

nucleation sites and therefore a different rate of nucleation.

0189888813 can cause an IROI‘BOSO in "18 CI'USIOI 8128 01 the COCOS,

therefore decreasing the density of potential dolomite nucleation sites.
Samples having undergone aggrading neomorphism and/or coarsely
crystalline cementation should therefore result in a substrate with a lower

density of dolomite nuclei per unit volume.

Substrate Surface Area Affects

(Examples)

There are few published analyses where the distribution of dolomite,
dolomite paragenesis, original lithologles, solution chemistry and porosity
are adequately described. The following examples of dolomite porosity
evolution are consistent with our ideas concerning the influence of
substrate surface area on dolomite porosity. In these studies, however, it
is not possible to assess what affects differences in original mineralogy
and/or changes in fluid chemistry had on the dolomite porosity
development.

Jodry (1969) examined dolomites from a reef and associated
carbonates (Silurian, Michigan Basin). He chronicled the evolution of the
porosity for a variety of original rock textures from 100 percent limestone
to 100 percent dolomite. He concluded that the 'rocks with the least
porosity development were micrite', with the resultant dolomites
exhibiting an average porosity of 1.8 percent. Dolomitized micrltes also

had a small dolomite crystal size in comparison to the dolomitized

20

reef -f ramework. He observed that very large crystals of dolomite and
higher porosity (12.5 percent) were apparent in dolomites that had replaced
reef -framework material. This porosity development is consistent with a
substrate surface area control on porosity.

Barrett (1986) suggested that the replacement dolomite of the
Smackover Formation (southern Alabama) and its porosity distribution,
were a function of changes in the magnesium flux and anisotropies in the
carbonate fabric. The effect of the carbonate fabric on porosity is
observed in the completely dolomitized areas. In these areas, the low
porosity dolomites were correlated with changes to finer-grained
carbonate substrates.

Armstrong (1970) examined dolomites from the Lisburne Group
(Alaska). In these dolomites, finer crystalline dolomite replaced the
carbonate mudstones and more coarsely crystal line dolomites replaced
packstones and grainstones. Abundant visible porosity was noted in the
dolomitized packstones and grainstones.

Fisher and Rodda (1969) examined dolomites from the Edwards
Formation (Texas). They suggest that there are two types of dolomite
present, 1)a stratal dolomite - which is fine grained, has a uniform grain
size, a tightly knit fabric, is very slightly porous and permeable, and is
associated with carbonate mudstones, and 2) massive dolomite - fine to
coarse grained, has a variable rhomb size, a loosely knit fabric, is
moderately porous and permeable, and is associated with fossiliferous
carbonate grainstones. The stratal dolomite is inferred to have formed
prior to lithification and the massive dolomite formed after lithification.
The differences in porosity between the two dolomite types may be due to
the initial differences in substrate surface area, and to the lithification

21

process which would increase the crystal size of the substrate. The
variable rhomb sizes in the massive dolomite may have resulted from; I)
coarsely crystalline dolomite replacing the coarsely crystalline fossils
that had been lithified and 2) the finely crystalline rhombs (similar in size

to the stratal dolomites) replacing the mud matrix.

Ranges of Solution Chemistry in Nature

Dolomitizing solutions consisting of a wide range of fluid
compositions have been inferred by many researchers. The differences in
the saturation indexes (lAP/Ksp) of these solutions with respect to
dolomite is examined below.

Dolomitizing solutions produced by various mixing proportions of fresh
water and seawater have been invoked by Badiozamani (1973); l'Ianshaw,
Back and Dieke (I971); Choquette and Steinen (1980); Sears and Lucia
(1980); Ward and Halley (1985) and many others. The saturation index of
this mixture using Badiozamani‘s (1973) data is approximately 14 (T able
2). Seawater has been suggested as a dolomitizing solution by Seller
(1984a); Mullins et al (1985); Carballo and Land (1987), Aharon et al (1987)
and others. Seawater, with a saturation index of approximately 2990
(T able 2) with respect to dolomite is more highly supersaturated with
respect to dolomite than mixed waters.

Ref luxing brines (which can have a wide range of compositions) have
also been suggested as a possible dolomitizing solution (Adams and Rhodes,
1960; Deff eyes et al, 1965; Sears and Lucia, 1980; Clement, 1985; Rosen et
al, 1988). The saturation index of the Dead Sea brine (Krumgalz et al,

22

TABLE 2. THE SATURATION INDEXES OF THE SOLUTIONS WITH RESPECT TO
ST OICHOHET RIC DOLOI'IITE

 

 

Solution Type Saturation Index
Mixed Water’ (.33 x 10'
Sea Water“ 2.99 x 103
BrineM 4.85 x 105

 

* mmitisn)nucusistsorsssrmrmmau5$mvm.

” Mmmemmmusanwtisee)mmmmmm
Drenr(1982).

“* Deedmbrimqproxinatehj imitheoonoentntionofseavatermlzuflfllerofim).

23

1982) which has been used for flux modelling is approximately 4.9 x 105
(Table 2). Although this particluar brine has a higher saturation index than
sea water, not all evaporatively concentrated waters are even
supersaturated with respect to dolomite (Gueddari, 1983).

The evaporitively concentrated brine has the highest saturation index
of the waters examined and the mixed water solution has the lowest. The
higher saturation index in the brine would lead to a greater nucleation rate
for the brine dolomites. The lower saturation index in the mixed water

would lead to a lower nucleation rate for the mixed water dolomites.

Solution Chemistry Affects
(Examples)

There are two examples where the porosity of the dolomites appears
to be directly related to the saturation state of the dolomitizing solution
(Longman et al, 1983 and Barrett, 1986). These studies however, do not
attempt to correlate the stratigraphic distribution of the dolomite with
changes in dolomite crystal chemistry, nor do they provide enough detailed
information to assess the role that changes in substrate surface area may
have in the porosity evolution of the dolomite.

Longman et al (1983) examined dolomites from the Red River
Formation (Ordovician, Williston Basin) detailing the dolomite distribution
and porosity trends (Figure 3). They showed dolomite pods under the C
anhydri to which crosscut bedding planes. The abundance of dolomite
decreased around the periphery of the pods, suggesting a decrease in the
Mg/Ca ratio of the dolomitizing solution. Near the source of the

24!

 

 

 
     
     

 

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11.15 llmeetone bed hectare bed
80-4 Ponmnrdokunue
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Figure 3. Schematic cross-section of the Red River Dolomites, Montana
(Longman et a1, 1983). Note the dolomite near the inferred source of the
dolomitizing solution (break in the C anhydrite) is cryptocrystalline and
has a low porosity. Dolomites farther away from the source of the solution
are more coarsely crystalline and have a higher porosity.

25

dolomitizing solution (an inferred break in the C anhydrite), the dolomite is
finely crystalline with low porosity. Near the limestone-dolomite contact,
the dolomite is more coarsely crystalline and porous. These dolomite and
porosity trends are interpreted to be related to decreases in the saturation
state of the solution with respect to dolomite.

Barret (1986) examined dolomites from the Smackover Formation
(Jurassic, Alabama) detailing the dolomite distribution and porosity trends.
It was suggested that areas of high magnesium flux determined the amount
and distibution of replacement dolomite and therefore its porosity. Areas
of high magnesium flux resulted in greater dolomite replacement, the
f orrnation of sucrosic dolomite and consequently greater porosity.

In this study constraints will be placed on the sediments to determine
what affects substrate surface area, solution chemistry or bicarbonate

flux would have on the evolution of dolomite porosity.

METHODS

Porosity analyses reported in this study come from three sources.
Core Labs (Mt Pleasant, Ml) performed whole core and plug porosity
analyses on selected samples from all sample locations. After these
values were obtained, thin section point count (500pts/slide) porosities
were determined on samples impregnated with blue epoxy or epoxy spiked
with a fluorescent dye. Fluorescent dyes were used to help determine
microporosity in the dolomites (Vanguas and Dravis, 1985). Regional trends
in the Trenton porosity were also determined from compensated neutron

porosity logs. A comparison of the dolomite whole-core porosity and the

26

compensated neutron log porosity was conducted in order to determine
regional trends in dolomite porosity (APPENDIX 1). Through the use of a
T-test, it was found that there was no significant difference in porosity
between the two methods.

Fossil percentages in the dolomites were obtained through point
counts of thin sections under fluorescent light (Dravis and Ilurewicz,
I985), diffused plane light, and plane light. Thin sections of the Trenton
Formation impregnated with fluorescent epoxy and viewed with a
fluorescence microscope reveal little or no microporosity. This indicates
that the thin section point counts are not biased by visible porosity.

Stoichiometry of the dolomites was determined by lt-ray diffraction
analyses (Graf and Goldsmith, 1956). Major and trace element analyses of
the limestones and dolomites were performed on a Perkin-Elmer '560 AAS.
The analytical and machine precision for strontium and manganese was
less than t 18, for zinc t 4 x, and for iron 1 208. All samples were
powdered to pass a 60 micron sieve and dissolved in glacial acetic acid
following the method outlined by Barber (1974).

Carbon and oxygen isotopic analyses were performed at the University
of Michigan Stable Isotope Laboratory. All samples were roasted under
vacuum at 380°C for one hour to remove volatile contaminants. Calcite and
dolomite samples were reacted in anhydrous phosphoric acid at 50°C, and

C02 was prepared in an extraction line coupled directly to the inlet of a VG
602E ratio mass spectrometer. Isotopic compositions were then converted

to P08 and corrected for ”0 according to Craig's (1957) procedure. No
fractionation correction was applied for the dolomite-phosphoric acid

reaction. Precision of the isotopic data is better than 0.10 per mil (0’00)

27

for both oxygen and carbon determinations. The precision of data is based
on a daily analysis of NBS 20.

OBSERVED AFFECTS OF SUBSTRATE
SURFACE AREA AND SATURATION STATE
OF THE SOLUTION ON DOLOMITE POROSITY

The Ordovician Trenton F orrnation
Jackson County, Michigan

The Trenton Formation (Ordovician, Michigan Basin) was studied
because it contains carbonate conglomerates and consequently provides
substrate textural variations within close proximity. In studying clasts of
strikingly different textures immediately adjacent to one another,
differences in solution chemistry and/or temperature can be dismissed.
The Trenton dolomites also provided an opportunity to look at how changes
in solution chemistry might affect dolomite porosity. These changes in
solution chemistry were inferred from the spatial distribution of the
dolomites and trace element gradients.

The Trenton Formation is composed of mudstones and wackestones
that were deposited in a deep subtidal environment (Wilson and Sungepta,
1985). In this study five cores from north central Jackson County were
examined (Figures 4 and 5). An isopach map of the Trenton Formation in the
Michigan Basin (Figure 6), indicates that the study area lies between what
are interpreted to have been two Trenton platforms. An isopach map of the

study area (Figure 7) reveals an anomalous local thickening of the Trenton,

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

 

29

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Formation in Michigan (from Wilson and Sengupta, 1985). Note the position

of the study area (star) between what are though to have been two Trenton
platforms.

31

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32

a feature that has also been seen in the Trenton from the nearby Northville
Trend (Mesher, 1980). A contour map of the top of the Trenton in the study
area (Figure 8), shows that there is no discemable local dip direction. The
regional dip in this area of the basin is to the north. Most of the Trenton
Formation in the study area is limestone, however, the cored intervals
contain partially and completely dolomitized strata.

The limestones are dominantly dark grey nodular mudstones and
wackestones. The allochems in the limestones are chiefly echinoid and
brachiopod fragments, with minor amounts of Iithoclasts, trilobites,
molluscs and pellets. Clasts are common in the upper part of the cores
(Figures 9 and 10). The clasts in this zone range in size from a few
millimeters to five centimeters. The clasts range in composition from
mudstones to crinoid-brachiopod-mollusc packstones, and generally
comprise about 65 percent of the rock volume where present. The crystal
sizes of the constituents in the limestone clasts vary widely (Table 3).
The fact that the clasts are rounded suggests that these clasts are detrital
and not collapse breccias. This is also supported by the presence of
rounded detrital quartz grains in the matrix between the clasts but not
within the clasts themselves (Figure I I). Molluscs in the limestones and in
the lithoclasts are often dissolved away with the molds being filled by an
equant blocky spar cement (Figure 12). Cementation occurred prior to the
deposition of the clasts. This is evidenced by cemented fractures in the
clasts which end abruptly at clast edges (Figure 13). These fractures do
not extend into the matrix and would not have survived transport if they
had not been previously cemented.

Oxygen isotopic compositions of whole-rock limestone analyses range

from -7.66 to -5.06 °l°° and average -6.39 °l°° P08. Carbon isotopic

33

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34

Figure 9. Cross-section A-A' showing the distribution of dolomite and
Iithoclasts. Gamma logs and neutron porosity logs for the wells are also
shown. Note the increase in neutron porosity in the horizons containing
dolomitized Iithoclasts. Also shown are the cored intervals of the wells.

35

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Figure 10. Cross-section B-B' showing the distribution of dolomite and
Iithoclasts. Gamma logs and neutron porosity logs for the wells are also
shown. Note the increase in neutron porosity in the horizons containing
dolomitized Iithoclasts. Also shown are the cored intervals of the wells.

37

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38

TABLE 3. VARIATION OF CRYSTAL SIZES (MICRONS) IN TRENT 0N LII‘ESTONE CLASTS

 

 

 

MUD MICROSDAR SDAR ALLOCI-IEMS
(I-‘I pl‘l) (‘1-50 pM) Short Long“ Short Long
RANGE IN SIZE 1-4 ‘1-17 1-40 6-50 6-810 12-2010
MEAN SIZE 2 7 13 30 1‘15 537
STANDARD DEVIATION 1 3 22 32 160 554
NUMBER OF SAMPLES 122 150 30 30 98 98
*Micrite and Microspar are considered to be spherical in sinpe. and Allochemsane

not and therefore they will have a maximum and minimum for be their short and
long dimension.

39

 

Figure 1 l. RDUT’IOBO OU8HZ grains In the matrix between dolomitized 818818
in the Trenton. Width of field : 3.8 mm

4O

 

Figure 12. Blocky low magnesium calcite cement filling fossil molds in the
Trenton limestone. Width of field : 0.6mm.

41

 

Figure 13. Cemented fracture in a limestone clast that is broken off at the
edge of the clast. Width of field = 3.8mm.

42

00111008111088 01 the WNOIB FOCK 11111881008 range from 1.01 10 ' 0.03 0’00

and average 0.15 °loo PDB (Table 4). The oxygen isotopic signature of these

limestones is approximately one per mil lighter that the isotopic
compositon of Ordovician sea water as estimated by Taylor and Sibley
(1986).

The minor and trace element compositions of the limestones are
presented in table 5. The strontium concentrations of the limestones range
from 562 to 235 ppm and average 405 ppm. The manganese concentrations
of the limestones range from 772 to 189 ppm and average 404 ppm. The
iron concentrations of the limestones range from 3606 to 1219 ppm and
average 2146 ppm. And finally, the zinc concentrations of the limestones
range from 10 to 3 ppm and average 6 ppm.

The amount of dolomite in the cores decreases gradually from west to
east and abruptly to the north and south (Figures 9 and 10). The dolomite is
planar to nonplanar, uniform to nonuniform (Sibley and Gregg, 1987) with
rhombs ranging from 25 to 1050 micrometers. The dolomite crystals
exhibit both undulose and nonundulose extinction. Inclusions in the
dolomites are so abundant that the dolomites appear light brown in plane
light. The original composition of the dolomitized limestones is difficult
to determine because most of the fossils in the dolomites are either ghosts
(particularly crinoids) or molds. However, the dolomitization crosscuts
lithologles, and therefore sections with undolomitlzed limestone clasts
probably represent the precursors to the dolomitized clasts. In the
horizons that have not been dolomitized, clasts range in composition from
mudstones to packstones (Figure 14). In order to decipher the original
texture of the dolomites, diffuse plane light and fluorescence microscopy

43

TABLE 4. OXVGEN AND CARBON ISOT OPE ANALYSES OF THE TRENTON
LINESTONES

 

 

 

88111018 1.11.1'101099 5 1801308 5 1381308
("’00) Woo)
T LUCK 2' 12
C182 Limestone ‘5. 17 -043
C282 ' -847 "031
C281 1 ' '880 '0.09
C389 ' '6.42 0.41
' -875 0.32
T Hankerd
C1815 ' '7.43 0.84
T F8181
4939 ' ’786 1.01
142'- 14 ' ‘506 '0.03
4945 ' '5.99 -033
118811 = -6.39 0.15

SD. = 0.89 0.52

44

TABLE 5. MINOR AND TRACE ELEMENT CONCENTRATIONS (PPM) OF THE
TRENTON FORMATION LIMESTONES.

 

 

Sample Strontium Manganese Iron Zinc
T Luck 2-12

C182 255 772 2831 3

C189 a 326 379 1219 9

b 559 386 1440 4

C289 a 552 470 I 767 12

b 302 456 2006 3

C2811 408 356 1950 10

C382 418 238 1312 9

C3813 485 406 2439 4
T Faist 1-12

C182 489 658 3606 7

C183 525 489 2304 4

C187 375 380 2363 4

C1812 235 282 2381 7

C1815 312 263 I456 _ 5

C1816 377 387 2771 3

C1818 300 189 2387 4

9.1.82.0 56.2 3.5.8 2.1.LL LQ

Mean = 405 404 2146 6

SD. = III 148 632 3

45

Trenton Limestone

Number of
Clasts

   

5 15 25 35 45 55 65 75 85 90
Percent Allochems

Figure 14. Composition of the Trenton Formation limestone clasts. The
percent of allochems in the limestone versus the number of clasts counted
with this percentage of allochems. The limestones range in composition
from mudstones to packstones.

46

were used to enhance fossil fragments (Figure 15 a and b). Areas that were
originally allochems often appear as grey inclusions in diffuse plane light.
In the dolomitized conglomerates, the dolomite which replaced the matrix
between clasts is finer than the dolomite which replaced the clasts (Figure
16 and table 6). Within the dolomitized clasts, the fossils are replaced by
rhombohedra that are larger than those in the adjacent matrix (Figure 17).
In areas of partial dolomitization in the conglomeratic zone, dolomite
preferentially replaced the matrix within the clasts and not allochems.
There is a small amount of ferroan dolomite cement which postdates the
initial and major dolomitization event (Figure 18). There is also a minor
amount of late sulphide mineralization that is seen in the dolomites. Minor
stylolitization is observed between some dolomite clasts.

The stable isotopic compositions of the Trenton dolomites are

presented in tables 7 and 8. The oxygen isotopes of the totally dolomitized

cores range from -6.49 to -10.42 °l°° and average -7.84 °loo PDB (Table 7).

The oxygen isotopes of the dolomites in the partially dolomitized cores
range from -6.07 to -7.63 °loo and average -6.57 °/00 PDB (Table 8).

The carbon isotopes of the totally dolomitized cores ranges from 0.58

to -0.67 °loo and average -0.02 °/00 PDB. The carbon isotopes of the
dolomites in the partially dolomitized cores range from 1.01 to -0.43 °loo

and average 0.15 °loo PDB. These values are similar to limestone values for

the Ordovocian (Veizer and Hoefs, 1976).

The dolomites are only slightly calcium rich with a mean value of 51
mole percent calcium (IIRD). The trace element concentrations of the
dolomites are presented in tables 9 and 10. The strontium concentration of

the dolomites ranges from 319 ppm to 96 ppm and averages 182 ppm.

47

Figure 15a. Trenton dolomite observed In plane light. Replaced allochems
appear as faint ghosts.

Figure 15b. Trenton dolomite as viewed in blue-violet light. Note the ease
with which allochems are observed.

48

 

Frgure 15a

 

 

Figure 15b

49

 

 

 

 

 

 

neuron
160
no [0
120
O
b
100
Hunter of .' ”5T8
Cryelele o o. MATRIX
so 0
b 0 O
40 i ..
O \I‘O’t I\.
20 9;) a
4.. to Page. 0. 0’.
00—49910 0 o-fiese «Mn-oe—

o 200 400 600 300 1000 1200
Mlcromeiere

Figure 16. Comparison of the crystal sizes between the dolomitized clasts
and the dolomitized matrix between the clasts in the Trenton Formation.

50

TABLE 6. DOLOMITE CRYSTAL SIZE DISTRIBUTION IN CLASTS FROM THE TRENT ON
FORMATION.

 

 

SIZE LUCK 1-12 LUCK 1-12 LUCK 1-12 LUCK 1-12 LUCK 1-12 TOTAL
(MICRONS) C1 Bl C1 BB CS 82 CS BS CS 812
1 -25 1 0 0 0 0 1
26-50 1 1 2 0 2 6
51 -75 1 2 3 1 4 1 1
76-100 2 4 4 2 3 1 5
1 01 -125 6 3 5 2 6 22
1 26—1 50 3 2 3 2 3 1 3
1 51 -1 75 4 8 6 6 5 29
1 76-200 20 4 7 7 4 42
201 -225 9 5 4 2 6 26
226-250 7 3 9 9 12 40
251 -275 7 2 3 5 2 1 9
276-300 20 7 5 8 14 54
301 -325 3 8 O 4 4 1 9
326-350 3 1 1 0 4 1 19
351 -375 1 4 13 7 1 0 9 53
376-400 3 0 1 3 2 9
401 -425 8 3 3 9 7 30
426-450 7 0 2 5 0 1 4
451 -475 1 2 0 0 1 4
476-500 1 0 1 1 3 5 20
501 525 3 0 0 3 3 9
526-550 4 1 1 4 2 1 2
551 -575 3 2 0 2 0 7
576-600 2 0 0 2 O 4
601-625 2 3 1 1 2 ' 9
626-650 1 0 0 1 0 2
651 -675 2 1 1 1 0 5
676-700 0 0 0 0 O 0
701 ~725 1 0 0 0 0 1
726-750 2 1 0 1 2 6
751 -775 0 1 1 1 1 4
776-800 4 2 1 1 8
801 -825 0 0 0 0 0
826-850 0 0 0 0 0
851 -875 0 1 0 0 1
876-900 2 O 1 3
901 -925 0 0 0
926-950 1 0 1
951 -975 0 0
976-1 000 O 0
1 001 -1025 0 0
1 026-1050 1 1

 

Total 51 9

51

TABLE 6 coni'd. DOLOMITE CRYSTAL SIZE DISTRIBUTION IN THE MATRIX FROM
THE TRENTON FORMATION.

 

 

SIZE LUCK 1-12 LUCK 1 -12 LUCK 1-12 LUCK 1-12 LUCK 1 -12 TOTAL
(MICRONS) C1 B1 C1 BB C3 82 C3 BS C3812
1 -25 12 3 7 1 1 4 37
26-50 34 20 36 28 22 140
51-75 36 25 19 14 24 118
76-100 33 23 18 1 1 23 108
1 01 -125 23 1 6 5 9 1 0 63
1 26-1 50 25 3 2 7 1 2 49
1 51-1 75 3 0 4 5 5 17
1 76-200 1 0 0 1 2 4 1 7
201 -225 7 1 0 0 1 9
226-250 3 0 3 0 1 7
251-275 2 0 0 0 O 2
276-300 0 1 2 1 0 4
301-325 1 1 1 1 1 4
326-350 0 0 0 0 0
351-375 1 0 0 0 1
376-400 0 O 0 0 0
401 -425 2 1 0 0 3
426-450 0 0 1 1
451 -475 0 0 0 0
476-500 2 1 1 4
501-525 1 1

 

Total 585

 

Figure 17. Large dolomite crystals replacing the allochems (A) and smaller
dolomite crystals replacing the adjacent matrix (8). Width of field :
3.8mm.

53

 

Figure 18. Late iron-rich dolomite cement In the Trenton Formation
dolomites. Width of field : 3.8mm.

54

TABLE 7. OXYGEN AND CARBON ISOTOPE ANALYSES OF THE TRENTON
DOLOMITES.

 

 

Sample Li thol ogy a ‘ BOPDB 6‘ 35P08
Woo) Woo)
T LUCK 1‘ 12
C181 100 3 0010171118 '7.71 “0.08
C 185 ' '81 1 0.28
C 188 ' '7.21 -067
' -849 0.50
' -730 ‘0.09
C281 ' ‘168 ‘027
' ‘722 “0.09
C382 " '750 0.24
' '7.71 0.05
C3812 ' -8.35 0.14
T LUCK 2’ 12
C283 ' -888 0.25
C288 ' '8.1 1 -028
' “'85? -053
C286 ' -753 0.04
C287 ' '729 0.03
C287 ' -732 0.12
T F8181
4898 ' "8.05 -028
4898 " '835 ”0.23
4899 .. '891 -008
$299 1 21.11.52 9.58
"8811 = “7.84 '0.02
5.0. = 0.85 0.31

55

TABLE 6. OXYGEN AND CARBON ISOTOPE ANALYSES OF DOLOMITES
FROM THE PARTIALLY DOLOMITIZED CORES, TRENTON FORMATION.

 

 

 

Sample Li thologg 6 IBOPDB s ISCPDB
(0’00) (01”)
T Hankerd 1" 17
C 182 P011101 0010171118 ’895 '0.01
C 183 ' '7.83 '0.12
C 188 ' “8.20 “0. 14
T Harmon LUCK 1'1
4930 ' '8.07 0.09
4941 ' “8.21 0.51
4950 " '8.38 0.37
Mean = “8.57 0.12
5.0. = 0.80 0.28

56

TABLE 9. STRONTIUM CONCENTRATION (PPM) OF THE TRENT ON
FORMATION DOLOMITES.

 

 

Sample Lithologg Nonporous Nonporous
Clasts Portions of
PM
T Luck 1-12
C181 a Dolomite 130
b ' I34 129
c " 155
C185 a " 157
b " 127 122
C281 " 159
C385 a " 155
D " 140 139
c " 150 123
C3812 " 162 188
C3816 a " 181
b " I41
C3817 a " 202 139
b " I61
C3821 " I29
C582 " 144
TLuck 2-12
C283 a " 203
b " 102
C286 a " 135 125
a " 96
C287 a " 132 125
Mean :2 146 143
3.0. = 27 23

Table 9 (cont'd.)

57

 

 

Sample Lithologg Strontium
T Hankerd 1-17

4829 Dolomite 174
4832 " 235
4833 " 251
4834 " 203
4837 " 277
4839 ' 319
4841 " 275
4845 " 311
4849 " 196
4854 " 232
4855 " 205
4857 " 263
4863 " 202
4865 " 273
4867 ' 303
$821 1 2.56

Hean = 248

5.0. = 44

58

TABLE 10. MANGANESE, IRON AND ZINC CONCENTRATIONS (PPM) OF
THE TRENT ON FORMATION DOLOMITES.

 

 

Sample Lithologg Manganese Iron Zinc
T Luck 1—12

13181 a Dolomite 2079 28200 3

b " 2001 25057 2

c ' 1622 29329 5

C185 a " 1777 21015 5

b " 1791 21733 2

C385 a ' 1714 17778 2

b " 1708 18913 1

c " 1685 19090 2

C3812 " 1819 18188 3

C3817 a " 1518 19262 3

l) " 1589 18329 4

C3821 " 1557 27143 2

C582 ' 1768 36486 9
T Luck 2-12

C283 a " 1736 19634 4

b " 1850 18329 5

c " 1385 15676 6

d " 1489 16221 5

8286 a " 1298 25891 2

b " 1425 23056 2

c ' 2078 34918 0

d " 1636 18879 5

C287 a " 1268 22112 2

b " 1213 24563 2

c " 1626 22677 5

n " .1261 2123.6 6

Mean = 1637 22789 4

SD. = 241 5456 3

59

Table 10 (Cont'd)

T Hankerd 1-17

4829 1388 14480
4832 1458 14779
4833 952 8789
4834 1710 17297
4837 1012 9903
4839 1144 10750
4841 1019 8938
4845 981 9551
4849 988 9230
4854 883 5733
4855 1114 11515
4857 1980 18450
4883 884 8000
4885 1891 15445
4887 997 8273
__437 I 522 491.8

1145 10543

399 4714

Iumud—ummauwmuw-‘A

NM

60

Strontium concentrations in these dolomites are similar to the
concentrations in other Ordovician dolomites (Weber, 1964) but higher than
many other ancient dolomites (Table 1 l). The strontium concentration in
the dolomites is roughly half of the strontium concentration in the
limestones (Table 5). The manganese concentration of these dolomites
ranges from 2079 ppm to 572 ppm and averages 1445 ppm. These
concentrations are higher than other Ordovician dolomites (Weber, 1964)
and other ancient dolomites (Table 11) but are similar to the
concentrations of other Trenton dolomites (Taylor and Sibley, I986). The
iron concentration of these dolomites ranges from 36,486 ppm to 4978 ppm
and averages 18,140 ppm. These concentrations are higher than the iron
concentrations of other Ordovician dolomites (Weber, 1964) but is similar
to other Trenton dolomites (Taylor and Sibley, 1986). The zinc
concentrations of these dolomites ranges from 10 ppm to 1 ppm and
averages 4 ppm. These concentrations are lower than other Ordovician
dolomites (Weber, 1964).

The whole core porosity of the Trenton Formation carbonates is
presented in Appendix II. The average porosity of the limestones is 1.3
percent, the average porosity of the partial dolomites is 1.9 percent, and
the average porosity of the dolomites is 3.1 percent. Whole core porosities
of the clast rich horizons are 0.8 percent (n = 32) for the limestones and
3.5 percent (n = 92) for the dolomites. The porosity difference between
Individual limestones clasts is low compared to the porosity difference
between individual dolomite clasts (Figure 19). The porosity in the
limestones occurs generally as fossil moldic (50 percent of the total
porosity). The dolomite porosity values determined by point counting were

found by subtracting the late f erroan dolomite and anhydrite cements, both

61

TABLE 11. STABLE ISOTOPIC (PDB) AND TRACE ELEMNT CU‘IPOSITIONS (PPM) OF SOME
ANCIENT DOLOMITES

 

 

 

 

Location Sr Mn Fe Sufi/0061300!” Proposed Origin
Ordovician, Tennessee 24 24? Fresh& Mixed
(Churnet and Misra, 1981)

Ordovician-Siluianfievada 49 476 02.4 12.0 Mixed
(Dunham&01ssn, 1980)

Mieeieeippian, Illinois 166 12.4 42.0 Mixed
(Choquette& Steinen, 1960)

Ordovician, Wisconsin 37 - 1.0 -4.5 Mixed

(Oadi ozamani, 1973)

Ordovician, N.W.T., Canada 81 46 10.4 -3.3 Mixed

(Land at al, 1975)

Silurian, Michigan 1 14 13.6 -7.4 Refluxing
(Sears 8. Lucia, 1980) Brines
Devonian, Albetra, Canada 57 59 343 93.0 -S.0 Burial
(Mattes a. Mountjoy, 1980)

Oowland Basin, England 60 10.5 -6.3 Ourial

(Bawthorpe, 1987)

 

NUMBER G:
CLASTS

 

 

I LIMESTONE CLASTS
I DOLOMITE CLASTS

 

 

Figure 19. Comparison of porosity between limestone and dolomite clasts.
Note the large range in porosity between the individual dolomite clasts.

 

63

of which are very rare volumetrically comprising less than one percent of
the total volume. The porosity in the dolomites generally occurs as
intercrystalline (89 percent), but can locally be attributed directly to
fossil molds (11 percent) (Figure 20). Dolomite clasts of differing
porosity are in close proximity and are often in contact with each other
(Figure 21). Horizons with dolomitized clasts generally have higher
porosities (based on compensated neutron logs, Figures 9 and 10) than
dolomitized horizons without dolomitized clasts, and higher than the

limestone sections.

Relationship of the Trenton Dolomites

to Fractures

Taylor and Sibley (I986) characterized the Trenton Formation
dolomites of the Michigan Basin into three distinct types based on major
and minor element chemistry, oxygen isotope ratios, and rock texture.
These dolomite types are; 1) cap dolomite, 2) regional dolomite, and 3)
fracture related dolomite.

The dolomites in this study are the fracture related type. Looking at
the areal distribution of the dolomite, the dolomite abundance in the area
decreases from west to east and from north and south. Two fracture
identification logs were run in the study area and indicate that the Trenton
Formation is locally highly fractured. The trends of the fractures are 11 27°
W and N 21° E (Figure 22). The northwest-southeast fracture trend is
similar to the regional fracture trends of the Albion-Scipio and the
Northville oil fields (File, 1962; Shaw, 1975 and Mesher, 1980).

64

 

Figure 20. Fossil moldic porosity in the Trenton dolomites. Width of field
3.8mm.

65

 

Figure 21. Dolomite clasts with varying porosity in close proximity. The
view presented is in plane light so the porosity appears white. A low
porosity clast (A), a moderate porosity clast (B) and a high porosity clast
(C) are in close proximity.

66

Figure 22. Rose diagram depicting the fracture pattern in the Trenton as
derived from two fracture identification logs run in the study area. The
trends of the fracture set are N 27°W and N 21°F.

 

 

 

68

The anomalous thickening in the Trenton as seen from the isopach map
is similar to the pattern in the Northville Trend. This thickening is
suggested to be due to fault related differential subsidence in the area
(Mesher, 1980). Donald Ferguson (personal communication; Total Petroleum
Company) indicated that the Trenton was highly fractured locally, and that
there was probably some type of horst and graben structure in the area.

These assessments were based on proprietary geophysical information.

The oxygen isotopes of these dolomites average -7.84 °loo P08. The

oxygen isotopes of these dolomites were compared to dolomites in the area
that were studied by Taylor (Table 12). The isotopic ratios in this study
are not significantly different from the Albion-Scipio or the Northville
fracture related dolomites. Both of these areas are considered to be high
temperature, fracture-related dolomites (Taylor and Sibley, 1986). The
oxygen isotopic values in this study are also, however, not significantly
different from the cap dolomites studied by Taylor (1982), which are
considered to be lower temperature dolomites. This may be explained by
either of the following; 1) that the dolomites in this study formed from a
solution that had a relatively low temperature, 2) or that there was a
lower water/rock ratio in these dolomites which affected the '80/ I°O
equilbrium ratios of these dolomites. With a lower water/rock ratio, a
greater percentage of the dolomites' isotopic signature would be inherited
from the precursor limestone, which is heavier than the associated
dolomites (Table 4).

Manganese concentrations in these dolomites are consistent with the
values of the Trenton fracture related dolomites, and are significantly
lower than the cap dolomites (Table 13).

iron concentrations in these dolomites are significantly higher than

69

TABLE 12. COMPARISON OF OXYGEN ISOTOPIC SIGNATURES BETWEEN

 

 

 

DOLOMITE TYPES
This Study Fracture Related Cap
(Albion-Scipio) (Northville)

l -7.7 -8.6 -7.9 -6.7
2 -8.1 -7.4 -7.3 -7.7
3 -7.2 -8.4 -8. 1 -7.6
4 -6.5 -7.8 -8.5 -8.4
5 -7.3 -9.4 -e.o -8.0
6 -7.7 -8.2 -7.4 -7.2
7 -7.2 -6.5 -8.6 -7.6
8 -7.5 -9.1 -9.0 -7.6
9 -7.7 -9. I -9.0 -7.8

10 -8.4 -6.9 -9.0 -7.8

1 1 -6.9 -9.1 -9.2 -7.0

12 -8.1 -9.7

13 -8.6 -8.2

14 -7.6 -8.8

15 ~73 -9.1

16 -7.3 -8.2

17 -8.1 -8.1

18 -8.4

19 -8.9

20 10.4

Mean = -7.8 -8.4 -8.4 -7.6

5.0. = 0.8 0.9 0.7 0.5

70

TABLE 13. COMPARISON OF MANGANESE CONCENTRATIONS (PPM) BET WEEN
TRENTON DOLOMITE TYPES.

 

Average Standard Number of
Deviation Samples
This Study 1384 414 44
Fracture Dolomites * 1477 520 33
Can Dolomites * 3786 1323 40

 

’Taylor and Sibley, i986

71

the values of the Trenton fracture related dolomites (although the ranges in
values overlap), and are significantly lower than the cap dolomites (Table
14).

Textural characteristics of these dolomites are similar to those
described by Taylor (1982) for the fracture related dolomites, and to those
described by Mattes and Mount joy (1980) for epigenetic dolomites.

These dolomites are generally coarse grained and nonplanar. Fossils are
generally not well preserved, and minor saddle dolomite is present. The
dolomites in the western portion of the study area have a higher percentage
of dolomite crystals with nonplanar boundaries (Table 15). The higher
percentage of nonplanar boundaries may indicate precipitation at an

elevated temperature (Gregg and Sibley, 1984).

Discussion of Substrate Surface Area Affects

In the Trenton Formation, one can constrain a number of factors
influencing the porosity of the dolomites, and subsequently show the effect
that the substrate surface area had on the resultant porosities. The
following parameters are constrained: I) the porosity variation observed is
not an inherited property of the parent limestone, 2) the porosity variation
is not a function of variance in solution chemistry or temperature, and 3)
that parent mineralogy has no verifiable control on the resultant dolomite
porosities.

First, the porosity difference observed between clasts in the
dolomites of the Trenton Formation was not inherited from the parent

limestone. High and low porosity clasts in the dolomite are often in direct

72

TABLE 14. COMPARISON OF IRON CONCENTRATIONS (PPM) BETWEEN
TRENTON DOLOMITE TYPES.

 

Average Standard Number of
Deviation Samples
This Study 18,142 7,666 41
Fracture Dolomites * 6,976 6,634 28
Cap Dolomites * 109,500 44,436 96

 

‘Taylor and Sibley, 1986

73

TABLE 15. POINT COLNTING RESllTS FOR DOLOMITE TEXTURAL COMPARISON

 

 

 

 

541-1911: PLANAR we TOTAL
(1111 11011 11111 11o11

(WEST)

TLuc1< 1-12 C3821 13 67 200
(2) (I I) (30) (57)

noon-12 use 17 63 201
(0) (I7) (5) (7B)

TLuck1-12 c362 14 aa 160
(0) (14) (0) (86)

Mean: 15 Mean = 85 TOTAL 581

(EAST)

7114mm 1-17c1aa 35 as 167
(0) (3S) (4) (61)

T 114mm 1-17 1:1 82 27 73 96
(0) (27) (1 1) (62)

THANKERDI-l'l 1:167 31 69 201
(0) (31) (0) (69)

THANKERDI-I" C1815 31 69 209
(o) (31) (1) (66)

Tammi-17 C183 26 72 206
(0) (2a) (1) (71) ‘

THANKERDl-I'} C1814 26 74 21s
(0) (26) (1) (73)

71141110290147 C1812 25 7s 71
(0) (25) (4) (71)

Tumsem-n c163 29 ' 71 231
(0) (29) (2) (69)

risen : 29 Mean = 71 TOTAL 1416

 

msemwmeMmmwm
m-demkmammm

74

contact with each other. Due to the immediate proximity of the clasts it is
therefore reasonable to assume that the clasts were subjected to the same
amount of bicarbonate flux. The existence of dolomite clasts with
approximately 608 porosity precludes the possibility that the observed
porosity in the dolomites was due to porosity in the original limestones.
The low average porosity of the dolomite (3 8 - APPENDIlt 11) indicates that
there has been at least 108 dolomite cement precipitated, according to the
volume shrinkage theory. if one assumes an average porosity of 108 for the
parent limestone, approximately 20: dolomite cement must have been
precipitated in order to produce the current average dolomite porosity of
38. Due to the large amount of dolomite cement in these clasts, the
porosity of the precursor limestone clasts would have to have been
unrealistically high. It is not likely that a clast of 80: porosity for
example, could have been transported and redeposited intact.

It is not likely that the limestone clasts had a wide difference in
porosity prior to dolomitization. The limestone clasts experienced a
porosity loss due to calcite cementation prior to dolomitization. In the
partial dolomites, dolomite crystals are observed replacing calcite cement
(Figure 23). This episode of cementation is also evidenced by the cemented
fractures in the limestone clasts. The fractures in the clasts end abruptly
at the clasts' edge and do not extend into the matrix (Figure 13). It is not
likely that these fractures would have survived transport if they had not
been previously cemented. If the limestones had an initially high porosity,
one would expect that in the partial dolomites there should be some
dolomite cement precipitated in the voids. This is not observed in the
partial dolomites. Also there is a minor amount of late evaporlte cementin

the dolomites. If there was abundant porosity in the limestones, one would

75

 

Figure 23. Nonplanar dolomite crystal (A) replacing calcite cement (8) in
the Trenton. Width of Field :1.5mm.

76

also expect to see evaporite cement in these voids. No evapori te cement is
observed in the limestones. The differences in porosity between individual
dolomite clasts was therefore not inherited from the limestone clasts, and
must then be the result of the dolomitization process.

Second, the close proximity of porous and nonporous dolomite clasts
(often in direct contact with each other) indicates that differences in
temperature and/or solution chemistry are not responsible for the observed
porosity differences. The porosity differences between clasts then, could
be the result of differences in substrate texture and/or mineralogy.

It is unlikely that the substrate had a significantly different
mineralogy. Flao et al (1987) have shown that dolomite replacing low
magnesium calcite and aragonite result in dolomites with significantly
different strontium contents. The dolomite that replaced low magnesium
calcite averages 210 ppm strontium and the dolomite that replaced
aragonite averages 660 ppm strontium. The similarity in the strontium
concentration observed between nonporous clasts and the nonporous
portions of porous clasts in these dolomites (Table 9) suggests that the
parent mineralogy of the clasts was similar. A difference in substrate
mineralogy is also not likely with the evidence of calcite cementation prior
to dolomitization. The metastable phases of aragonite and high magnesium
calcite are generally the first phases affected by diagenesis. It is likely
that if these metastable phases were in the sediment prior to the calcite
cementation, they would have been converted to more stable low
magnesium calcite.

There are several observations that are consistent with the hypothesis
that an original difference in substrate texture caused the differences in

porosity observed. Differences in the substrate texture between clasts

77

(Figure 14), and crystal sizes between constituents (Table 3), have been
demonstrated. llhombohedra size variation within and between the
dolomite clasts suggests that the texture of the substrate did have an
effect on dolomite nucleation density and therefore its resultant texture.
For example, where decernable, the dolomite rhombohedra that replaced
fossils are larger (due to low nucleation density) than the rhombs that
replaced the adjacent matrix (due to high nucleation density) (Figure 17).
Also, the highly porous clasts in the dolomite have large dolomite crystal
sizes and the low porosity clasts have relatively smaller dolomite
crystals. This demonstrates the direct relationship that substrate texture
had on the resultant dolomite texture/porosity. Dolomitized clasts also
directly show some fossil moldic porosity. This also demonstrates the
direct relationship that differences in substrate texture had on the
resultant dolomite porosity.

These observations are consistent with the hypothesis of substrate
influence on texture and therefore porosity. According to this hypothesis,
the variability of the constituents in the clasts resulted in the porosity
differences observed between dolomite clasts. Specifically, the porosities
in the dolomites are directly related to the amount of coarsely crystalline
allochems, cement and/or neospar in the original limestone. This is based
on the assumption that the coarsely crystalline portions of the substrate
would resist dolomitization. This is evidenced in the partial dolomites
where finely crystalline mud is replaced by dolomite and the coarsely
crystalline allochems are not. The coarsely crystalline portions of the
substrate were therefore unreactive to the onset of dolomitization. The
resistence to dolomitization by coarsely crystalline portions of the

substrate is also evidenced in the totally dolomitized rocks by the

78

existence of fossil moldic porosity. These coarsely crystalline portions of
the substrate resisted dolomitization and occupied space during much of
the dolomite replacement and cementation. These coarsely crystalline
portions of the substrate were then later dissolved out, leaving a
high-porosity clast. The coarsely crystalline fossil and/or spar-rich
limestone clasts (low surface area) provided fewer nucleation sites than
the finely crystalline mud rich clasts (high surface area) for the
dolomitization episode. This resulted in the clasts with the higher
concentration of coarsely crystalline fragments becoming the dolomite
clasts with the higher porosity. Dense nucleation and interference of
dolomite rhombs occurred in substrates that had a high surface area

(mudstones) resulting in the lower porosity clasts.

Discussion of Solution Chemistry Affects

The distribution of the dolomite in this section of the Trenton
suggests that the solution flowed in a general easterly direction, because
the dolomite decreases in abundance to the east (Figure 9). Trace elements
and oxygen isotopes of the Trenton dolomites were examined in order to
confirm the flow direction of the dolomitizing fluid.

Veizer (1983) and Parker et al ( 1985) suggested that changes in the
gradients of trace elements may indicate the direction of fluid f low during
diagenesis. Land, Salem and Morrow (1975), Land, (1980), Gregg (1988) and
Machel (1988) suggest that trends in trace elements can be used to
determine the fluid flow direction of dolomitizing solutions, if the system

is not completely open. Trace elements can be used because during

79

dolomitization elements will fractionate between the dolomite and the
aqueous phase. Elements with a distribution coefficient of less than one
will fractionate to the aqueous phase. The actual amount of a phase that is
incorporated into a crystal is a function of the distribution coefficient, and
the total concentration of the species in solution (Veizer, 1983). In the
case of a distribution coefficient less than one, as the solution evolves, the
concentration of the species in solution will increase. The crystals
forming in the downf low direction should therefore have a higher
concentration of this species in the crystal lattice. This scheme assumes a
constant distribution coefficient. Although changes in temperature,
solution chemsitry and many other factors may affect distribution
coefficients, the magnitude of these effects on the distribution
coefficients for dolomites are not known. In this study the distribution
coefficient are assumed to be constant. Elements with a distribution
coefficient greater than one will tend to be incorporated into growing
crystals. The concentration of this species in solution will therefore
decrease in the downf low direction. The dolomite crystals in the downf low
direction will therefore have a lower concentration of this species. I

The actual distribution coefficient of Sr2+ is still under debate (Land,
1980). It is agreed however, that it is less than one (Land, 1980, Veizer,
1983, Burns and Baker, 1987). The dolomite crystals in the downf low
direction should therefore become enriched in 59*. The Sr2"
concentrations of the dolomites in the Trenton Formation average 143 ppm
(sd = 26, n = 29) for the Total Luck wells (west) and 245 ppm (sd = 43, n =
20) for the Total Hankerd well (east). A W-test was conducted on the 512+
concentrations from the Total Luck wells, and it was determined that these

values were from a normally distributed population (Appendix III). A

80

standard T-test was then conducted on the two populations. The Sr”
concentrations from the Total Luck wells and the Total Hankerd wells are
significantly different at the 98: confidence interval (Appendix III). This
type of trend in the Sr1H concentration is consistent with the stratigraphic
distribution suggesting an easterly fluid flow.

Much less is known about the distribution coefficients of manganese,
iron and zinc in dolomites. However, the distribution coefficients of all of
these elements are considered to be greater than one (Land, 1980, Veizer,
1983). The Mn2+ concentrations in the dolomites of the Trenton Formation
average 1637 ppm (sd = 241, n = 24) for the Total Luck wells (west) and
1 145 ppm (sd = 399, n = 16) for the Total Hankerd well (east). A W-test
was conducted on the Mn” concentrations from the Total Luck wells, and it
was determined that these values were from a normally distributed
population (Appendix III). A standard T-test was then conducted on the two
populations. The Mn” concentrations from the Total Luck wells and the
Total Hankerd wells are significantly different at the 95! confidence
interval (Appendix III). This trend in Mn2+ concentration is also consistent
with an easterly fluid flow.

The Fe2+ concentrations in the dolomites of the Trenton Formation
average 22,789 ppm (sd = 5456, n = 25) for the Total Luck wells (west) and
10,543 ppm (sd = 4714, n = 16) for the Total Hankerd well (east). A
W-test was conducted on the F e2+ concentrations from the Total Luck
wells, and it was determined that these values were from a normally
distributed population (Appendix III). A standard T-test was then
conducted on the two populations. The Fe2+ concentrations from the Total
Luck wells and the Total Hankerd wells are significantly different at the
98: confidence interval (Appendix III). This trend in Fe2" concentration is

81

also consistent with an easterly fluid flow.

The Zn2+ concentrations in the dolomites of the Trenton Formation
average 4 ppm. The lack of any trend in Zn” concentration may be due to
the low absolute concentration of zinc in the solution.

Oxygen isotopic composition of the dolomites could also be an
indicator of fluid flow in these dolomites. If these dolomites were
precipitated from a dolomitizing solution with an elevated temperature,
then as the solution moved away from its source it should cool and produce
a dolomite with a less depleted isotopic signature. Consequently, the trend
that should be observed in the Trenton dolomites is an enrichment in the
oxygen isotopic signature of the dolomites in an easterly direction. The
oxygen isotopic composition of the Trenton dolomites from the Total Luck
2-12 well (west) to the Total Hankerd well (east) are presented in figure
24. There is a general increase in the oxygen isotopic signature from west
to east. This data also supports an easterly direction for fluid flow. It is
possible however, that this trend simply reflects the effects of a lower
water/rock ratio. in this case, the oxygen isotopic composition of the
substrate would have an increased role in determining the final isotopic
composition of the dolomite.

There is also a change in dolomite texture which is consistent with
fluid flow in an easterly direction. Gregg and Sibley (1984) determined
that the temperature of the dolomitizing fluid had an affect on the texture
of the dolomite. Specifically, a solution with a high temperature will
result in a higher proportion of nonplanar dolomite. The change in texture
observed is a statistically higher proportion of planar dolomite (Cl 952) in
the eastern portion of the study area. This change in texture is consistent

with fluid flow in an easterly direction, with the temperature of the

82

 

 

 

 

 

 

-s.o--
O
6.5" O
0
'7.0" 9
3 0 TRENTONDOLOMTTE
18
6 O 7.50 6 . ° TrendofTRENTON
e DOLOMITE
-e.0«-
O O
O
“8.5+ .
-9_o : : : : .
0 1 2 3 4 5

DISTANCE (km)

Figure 24. Oxygen isotopic signature of the dolomites versus distance from
west to east for the Trenton dolomites. The trend is towards a heavier
oxygen isotopic signature of the dolomites towards the east.

83

solution decreasing away from the source.

Evidence that the dolomitizing solution flowed in an easterly direction
is: I) the stratigraphic distribution of the dolomites, 2) the changes in the
trace element compositions of the dolomites, 3) the trend in the oxygen
isotopic composition and 4) the difference in the texture of the dolomites.
Therefore the saturation state of the solution with respect to dolomite
should decrease in this direction. Recall from above that solutions with a
higher saturation state should result in a higher density of dolomite nuclei,
resulting in a lower dolomite porosity.

One can look at the porosity of the dolomites across the field to
determine if the change in saturation state that dolomitized these rocks
caused a difference in porosity. Compensated neutron log porosities of the
Total Hicks-Martin 1-1 1 well in the west and the Total Luck 2-12 well in
the east were compared (APPENDIX IV). There is no statistical difference
in the porosities between the two wells. A comparison of the whole core
porosities between the conglomeri tic portions of the Emil Faist 2-12
(west) and the Total Luck 1-12 (east) wells was also made (APPENDIX IV).
There was also no significant difference in the porosities between these
wells. This indicates that the change in fluid chemistry that these
limestones experienced did not significantly influence the porosity of the
resultant dolomite. This is not to say however, that greater changes in

fluid chemistry could not have an influence on porosity.

Pliocene Sero Domi Formation

Aruba, Netherlands Antilles

Two outcrop locations of the Sero Domi Formation on the island of
Aruba were examined. The two outcrops investigated in this study (Figure
25) are steeply dipping forereef deposits. The outcrop at Boe Doei will be
examined first with regard to the affects of substrate surface area on
dolomite porosity. The outcrop at Flooi Hundu will be examined later with
regard to the bicarbonate flux model. The island of Aruba, Netherlands
Antilles is located approximately 30 kilometers north of the peninsula of
Paraguana', Venezuela. The upper Sero Domi Formation exposed on the
island of Aruba is composed of a series of forereef limestones and
dolomites (DeBuisonje, 1974). The rocks are predominantly coralline algal
grainstones and packstones (Sibley, 1980; 1982). The dolomites at both
locations are conforrnably overlain by limestones.

The rocks at Boe Doei dip approximately 11° to the north. The
limestones exposed are mainly wackestones. The major allochem in these
samples is pelagic f oraminif era (Signage/rips sp.) which can comprise up
to 44 percent of the rock volume. Minor amounts of coralline algae and
echinoderms are also present. There is little or no evidence of compaction
in these rocks. li-ray diffraction analysis indicates that the rock is low
magnesium calcite. There is a euhedral blade-shaped calcite cement in the
limestones which comprises up to 25 percent of the rock volume. The
porosity of these rocks averages 17.1 percent (Figure 26 and Table 16). The
permeabilities of these rocks vary widely from 1.5 to 9530 rod (Table 16).

Whole rock oxygen isotopic composition of the limestones ranges from
-3.76 to -2.20 °loo and averages -3.15 °l00 P08. Whole rock carbon isotopic

84

85

8:25 88.8562
.334 :6 8838 8:686... .58 Bow 2: .22 SE .3383 .m« 2.3:

  
   
  

(JuDNuzw>

O c2556... .660 80m

0 23:86.. :23..."

ate-.1 .OOC

<m3m<

86

 

 

 

 

Boe Doei
O T . Dolomite
O O
2 .. 0 ° Limestone
O
o I Partial
4 0° 0 Dolomite
0
- O I
Depth In 6 0 -
Meters l
O
.0
8 :1-
O . O
O
10 up . .
O
12 : : 3 : : c : 9%

 

 

5 1015 20 2530 35 40455
Porosity

Figure 26. Percent porosity versus depth for the carbonates from Boe Doei
(Aruba, NA).

87

TABLE 16. 906061711 AND PERMEABILITY or THE 80E DOEI CARBONATES (ARUBA,
NA)

 

 

 

SAMPLE DEPTH IN LITHOLOGY PERMEABILITY POROSITY METHOD
METERS (MILLIDARCIES)
802A 1.0 LIMESTONE 1.5 24.7 PLUG
8028 1.5 LIMESTONE 7.0 TS
803A 2.0 LIMESTONE 6484.0 24.8 PLUG
8038 2.5 LIMESTONE 11.6 TS
804A 3.0 LIMESTONE 7275.0 28.4 PLUG
8048 3.5 LIMESTONE 7.4 TS
807A 4.0 LIMESTONE 9530.0 23.3 PLUG
8D78 4.5 LIMESTONE 17.6 TS
809A 5.0 LIMESTONE 1.4 1 i .9 PLUG
801A 5.2 OS. LIMESTONE 4202.0 22.0 PLUG
8018 5.5 OS. LIMESTONE 18.2 TS
808A 6.0 OS. LIMESTONE 1615.0 20.1 PLUG
8088 6.5 OS, LIMESTONE 8.4 TS
805A 7.0 DOLOMITE 29128.0 45.8 PLUG
8058 7.5 DOLOMITE 39.6 TS
806A 8.0 DOLOMITE 8654.0 38.5 PLUG
8068 8.5 DOLOMITE 30.0 TS
8060 8.8 DOLOMITE 33.2 TS
8010A 9.0 DOLOMITE OFF SCALE 48.5 PLUG
80108 9.5 DOLOMITE 38.2 TS
801 00 9.8 DOLOMITE 35.9 TS
8811A 10.0 DOLOMITE 0.02 7.6 PLUG
801 18 10.5 DOLOMITE 7.2 TS
Limestone porosity average - 17.1 % Permeability average - 46481er
Standard deviation - 8.2% Standard deviation - 4390an
Partial dolomlte porosity average - 17.2% Permeability average - 2909md

Standard deviation - 6.1%

Dolomite porosity average - 32.5 %
Standard deviation - 14.3%

88

composition ranges from -8.42 to -5.51 °/0° and averages -7.05 °l°o P08

(Table 17).

The dolomites are planar, uniform dolomite rhombohedra with an
average crystal size of 230 micrometers. Inclusions of allochems and
calcite cement are rare, making the texture of the limestone that the
dolomite replaced difficult to decipher. There are patterns of inclusions
present however, that can be identified as replaced foraminif era.
Inclusions of low magnesium-calcite cement in the dolomite are similar in
morphology to the low magnesium-calcite cement in the overlying
limestone (Figure 27). This indicates that a period of cementation
preceded dolomitization in this area. Glauconite pellets are observed in the

dolomites, but comprise less than one percent of the total volume. The

dolomites are calcium-rich with a mean value of 5? mole percent tacos

(X80). The dolomite trace element data and the whole rock stable isotope
composition are presented in table 18. The strontium concentration of the
dolomites ranges from 27? to 140 ppm and averages 176 ppm. The iron
concentration of the dolomites ranges from 2663 to 1265 ppm and averages
1961 ppm. The zinc concentration of the dolomites ranges from 6? to 15

ppm and averages 37 ppm. Whole rock oxygen isotopes of the dolomites

range from 3.74 to 2.26 °/00 P08 and average 3.32 °/00 P08. whole rock
carbon isotopes of the dolomites range from -3.?5 to - 3.09 °l°o P08 and

average -3.24 0/00 P08. The porosity of the dolomites ranges from 7.2 to

48.5 percent, with an average porosity of 32.5 percent (Fig. 26 and Table
16). The porosity in the dolomites is intercrystalline. The permeabilities
of the dolomites vary widely from 0.02md to a permeability that went off
scale (Table 16).

TABLE 17. OXYGEN AND CARBON ISOT OPE ANALYSES OF THE SERO

DOHI LIMESTONES, (BDE DOEI).

89

 

 

Sample Lithology s IBOPDB 5 130,303
(0’00) (0,00)
Boe Doei
A Limestone -3.5 1 -7.42
B ' '220 “5.5 I
C .. ‘355 '7. 15
D " “3.07 “'6. I 0
E ' '3.35 "7.60
F " ’3.76 '8.42
G ' '27 I ”6.87
H " :19} 71.19
"88" = -115 ’7.05
5.0. = 0.5 I 0.90

90

 

Figure 27. Blade shaped low magnesium calcite cement crystal included in
a dolomite crystal from Boe Doei (Aruba, NA). Width of field = 0.6mm.

91

TABLE 18. STABLE ISOT OPE AND TRACE ELEMENT ANALYSES OF THE
SERO 001‘11 DOLOMITES (BOE DOEI).

 

 

Sample £150,303 £130,303 Sr (DDm) Fe (ppm) Zn (ppm)
Woo) Woo)
Boe Doei
5 2.26 -3.75 277 2029 59
163 1664 67
6 3.43 -3. 13 16B 1620 63
' 171 1954 55
6 3.50 -3. 18 165 1273 19
10 3.72 -3. 14 176 2595 26
169 2590 29
1 1 3.74 -3.16 140 2663 26
157 1265 15
1.1. 3.25 2.3.99 .151 12.7.3 1.5
mean = 3.32 “-3.24 176 1961 39
SD. = 0.55 0.25 37 554 21

92

The limestone-dolomite contact parallels bedding, and occurs over a
vertical distance of approximately one meter. Across the contact, the
number of dolomite rhombs decreases but the size of the dolomite rhombs
remains the same (average of 256 micrometers). The dolomites in this
zone also exhibit the same calcium rich zone as dolomites from the areas
of 100 percent dolomite (Figure 28). There were therefore fewer dolomite
nuclei in these limestones, but the rhombs grew at the same rate and at the
same time as the rhombs in the completely dolomitized section. The fact
that the dolomitization parallels bedding is also consistent with the idea
that the degree of dolomitization is related to the original sediment

texture or composition.

Discussion of Substrate Surface Area Affects

0n the island of Aruba, most of the dolomites from the Boe Doei
outcrop are very porous, but two have a very low porosity (Figure 26 and
Table 16). The difference in the porosity of the dolomites at Boe Doei can
also be explained due to a difference in the host rock texture. A number of
factors that influence dolomite porosity could be constratined,
specifically: I) that the porosity variation was not a function of
differences in solution chemistry or temperature, and 2) that parent
mineralogy has no verifiable control on the resultant dolomite porosities.

First, the similarity in oxygen isotopes and the trace elements of the
dolomites (Table 18) argues against a major difference in pore fluid
chemistry and temperature.

Second, differences in the strontium content of the porous and

93

 

 

Figure 28. Dolomite crystals from Boe Doei under crossed-polars. Note the
calcium-rich zone towards the outer edge of the crystals in the completely
dolomitizied section (A). The same high calcium zone is seen in the
dolomites from the partially dolomitized horizons (8). Width of field =
1.5mm.

94

nonporous dolomites are not significant enough to suggest that there was a
major difference in the amount of aragonite between the samples (09. Hao
et al, 198?).

Third, both the high and low porosity areas have dolomite crystals of
similar size. The lower porosity of the dolomites in this portion of the
section was therefore the result of a higher initial packing of dolomite
nuclei. The subsequent growth of the dolomite crystals with the more

dense nuclei resulted in a lower porosity dolomite.

BICABBONATE FLUX

The flux of bicarbonate affects the porosity of a dolomite in that it
controls the amount of material available for the precipitation of dolomite
cement. Systems having high flux rates of bicarbonate, are able to
precipitate greater quantities of cement.

To form a dolomite with a porosity less than or equal to the porosity
of the precursor limestone, the amount of dolomite cement that must be

precipitated is at least 13 percent of the solid volume of the limestone.

The limiting ion for thls precipitate is llkelu to be 111203. because 118

concentration in solution is much less than the concentration of magnesium
ions in most dolomitizing solutions.

The amount of cement that has been precipitated can be easily
determined from the porosities of limestones with the some original
texture as the dolomites. Once the amount of cement precipitated has been
determined, the amount of time required to precipitate this cement can be

calculated. To calculate the amount of time required to precipitate the

95

cement, the mass flux of the reactants must be modelled. The dolomite

cementation equation used for model ling is:
1192* + ca2+ + 2Hco3' ---- cangtco312 + 2H“ (5)

The mass flux equation is:
Mass Flux = dqc (6)

d = water density
q = Darcy velocity

c -.- molality of the reactant

The Darcy velocities used are from Simms (1984; table 19). If the mass
flux of the reactants, the thickness of the dolomitized unit (assuming
vertical flow), and the amount of cement to be precipitated are known, the
minimum time for the precipitation of the cement can be determined from

the following:

{2 moles H603
Timestyers)=( mole g It _{[ (Lao-bag )i._ll'o_(_l_-_m)11(t_hfikne_ss(m))1 (7)

(mass flux {moles HCOQ) )t(molar volume of Da - r73
\ biz-year I (mole 03}

 

Three different dolomitizing solutions were examined for use in the

bicarbonate flux model. These solutions are; 1) mixed fresh/seawater, 2)

96

seawater, and 3) evaporatively concentrated seawater. It is difficult to
find examples where both the chemical analyses of the solutions and their
activity coefficients are listed. The activities for modelling a mixed
water flux are obtained from Badiozamani (1973). The data are for a
mixture of 5 percent seawater with 95 percent freshwater. The activities
for modelling sea water were obtained from l'lillero and Schreiber (1982)
with the absolute concentrations of the ions from Drever (1982). The
activities of ions in brine solutions are difficult to determine and are
therefore often not reported. The activities and concentrations for
modelling brine reflux are from Krumgalz and Hillero (1982). The
concentration of the brine is approximately ten times the concentration of
sea water.

The amount of bicarbonate available for cementation was determined
by taking bicarbonate and Ca” out of solution down to calcite equilibrium.

Recall that the “local source theory' invokes the local dissolution of the

CaCD3 substrate for dolomitization to proceed. 8y staying above calcite

equilibrium, the amount of dolomite cement that can be precipitated from
solution can be determined because dissolution of the substrate has not
yet commenced.

The flux of bicarbonate for the various flow systems modelled is
presented in table 19. The highest flux rates are in the mixed water flow
process and the lowest flux rates are in the refluxing brines. The flux
rates for seawater circulation are those associated with mixing zones.
The high flux rates for mixed waters is due to the high Darcy velocity of
this system. This particular mixed water however is undersaturated with
respect to calcite and therefore cannot precipitate dolomite cement

according to this method. Neither of the two locations examined in this

97

TABLE 19. DARCY VELOCITIES AND MASS FLUXES

 

 

Flow Process Darcy Velocity" I'lass Flux
(meters/year) (moles/meter2 year)

Hc03'
Reflux 0.01 -- 0.1 8.0x10'2 - 8.0x10"'
Kohout 0.3 - 1.0 2.7x10" - 9.2x10"

Coastal Flows

Mixing zone 5.0 - 20.0 6.5x10° - 2.6x10'
Seawaun- <15 - :so 46m10" - 1276x10°
Circulation

 

*Simms ( 1284)

98

study however, are suggested to have been dolomitized by mixed water.
The flux model will be applied to two study areas. First, the Pliocene
carbonates on Aruba (Netherlands Antilles) are examined because they
provide an opportunity to compare limestones and dolomites that had
similar original textures. By comparing porosities of the limestones and
dolomites one can estimate the amount of bicarbonate that has been
imported to, or exported from, the system. Second, the l'liocene carbonates
of the Eniwetok Atoll are examined because constraints on the timing of
dolomitization and the type of dolomitizing solution are available from

previous work.

Pliocene Sero Domi Formation
Aruba, Netherlands Antilles

The rocks at 800i Hundu (Figure 25) were used for mass flux
modelling because the original texture and porosity of the limestones can
be deciphered. This allows one to determine the amount of bicarbonate
that was f luxed through the system to produce the observed porosity in the
dolomite. The limestones range in composition from wackestones to
packstones. The major allochems are coralline algae which comprise up to
48 percent of the rock volume, and benthic foraminifera which comprise up
to 30 percent of the rock volume. Minor constituents include echinoderms
and rarely coral fragments. 8qu mineralogy as determined by x-ray
diffraction is low magnesium calcite. There is a very finely crystalline
acicular calcite cement in the limestones. Supko (1971) and Chafetz et al

(1985) suggest that cement crystals of this morphology are indicative of

99

vadose cementation. The limestones near the top of the outcrop are the
most tightly cemented. Based on the crystal morphology and the proximity
to the top of the present outcrop, the calcite cement is probably modern
vadose cement. The porosity of the limestones varies from 6.8 to 32.2
percent, with an average porosity of 21.7 percent. The porosity in the

limestones is intercrystalline, intracrystalline and vuggy. The whole rock

oxygen isotopes of the limestones range from -5.25 to -3.93 °l00 and
average -4.64 °/0° PDB. The whole rock carbon isotope values range from

-9.92 to -6.53 °loo and average -8.00 °loo PDB (Table 20).

The dolomites at Root Hundu are planar nonuniform dolomites with
rhombs averaglng 13.8 micrometers in diameter. The coralline algae in the
rocks were mimlcally replaced whereas the foraminifera were generally
dissolved and left as molds (Figure 29 a, b and c). There are however some
forams that have been replaced by dolomite. In contrast to the dolomites
at Boe Doei, there is no evidence of predolomlttzatlon cementation in the
rocks from this location. illnor compaction fractures are observed in the
coralline algae at this site. This compaction occurred prior to the
dissolution of the foraminif era. This is evidenced by coralline algae that
have been deformed around what are currently foram molds (Figure 30).
There is some post-dolomitization calcite cementation In the dolomites
near the quartz diorlte basement rock. This cement is a blocky phreatlc
low-magnesium calcite cement. The dolomite 1s calcium-rich with a mean

value of 57 mole percent CaCD3 (XFID). The dolomite trace element data
and whole rock stable isotopic composition is presented in table 21. The

strontium concentration of the dolomites ranges from 375 to 197 ppm and

averages 259 ppm. The iron concentration of the dolomites ranges from

TABLE 20. OXYGEN AND CARBON ISOTOPE ANALYSES OF THE SERO

100

DOHI LIMESTONES, (ROOI HUNDU).

 

 

Sample Lithology s‘BoPDB 5‘3ch3
(Woo) (Woo)
R001 HUI’IOU
A Limestone -s.25 -9.92
B " ~s.09 45.53
c ' -4.49 -7.07
0 ' -3.93 -9.10
E " 21.52 1.59
rlean = -4.64 —s.00
3.0. -.- 0.54 1.44

101

 

Figure 29a. Typical limestone from Rooi Hundu. Note the presence of
coralline algae and benthic foraminifera. Width of field = 3.8mm.

102

 

Figure 29b. Dolomites from Rooi Hundu showing the mimic replacement of
the coralline algae. Width of field : 3.8mm.

103

 

Figure 29c. Foram molds in the dolomites from Rooi Hundu. Width of field
=38mm.

104

 

Figure 30. Coralline algae fractured around what is currently a foram mold
from Rooi Hundu. Width of field 1.5mm.

105

TABLE 21. STABLE ISOTOPE AND TRACE ELEHENT ANALYSES OF THE

SERO DOi’il DOLOMITES, (ROOI HUNDU).

 

 

Sample 51BOPDB 5 1 3CPDB Sr (130"!) Fe (091") Zn (DOM)
(0’00) (0’00)
R001 ”[11101]
A 3.06 ‘2. 12 375 620 21
293 916 2 1
B 3.17 ‘1.92 271 725 18
266 696 16
C 1.90 -351 232 776 10
250 35 1 20
D 2.20 ‘4.05 263 356 25
227 324 6
E 1.77 ’2.23 197 306 25
F 2.2.1. £3.12 1.9.2 3.1.9 2.1.
"8611 = 2.39 “2.99 259 579 19
3.0. = 0.60 1.01 52 266 6

106

918 to 308 ppm and averages 579 ppm. The zinc concentrations of the

dolomites range from 25 to 6 ppm and averages 19 ppm. Oxygen isotopes

of the dolomites range from 3.17 to 1.77 °loo and averages 2.39 °/00 P08.
Carbon isotopes of the dolomites range from -4.12 to - 1.92 °l00 and

average -2.99 °l00 PDB. The porosity of the dolomites ranges from 20.7 -

35.5 percent, with an average porosity of 28.6 percent (Figure 31 and
Table 22). Vug and fossil moldic porosity are dominant in the dolomites.
The permeabilities of the limestones range from 5.9 to 6217md.

The limestones conforrnably overlie the dolomites, but It is possible
that these sediments were deposited after the underlying sediments had
been dolomitized. This is postulated because the limestone-dolomite
contact is sharp, parallels bedding, and partially dolomitized rocks are
absent.

80th the limestones and dolomites are coralline algal packstones to
wackestones. The affect of dolomitization on the porosity of these rocks
can be estimated by subtracting the post-dolomitization calcite cements
In the dolomites and the calcite cement in the limestones. With this
correction, the porosity of the dolomites (ave. 28.1 percent) is less than
the porosity of the limestones (ave. 39.9 percent) (Figure 31 and Table 22).
The dolomitization of these rocks has therefore lowered porosity, despite
the fact that fossil molds were created. This indicates that there must be
some dolomite cementation. The perrneabilitles of the dolomites range
from 594 to 19,997 mi (Table 22).

107

 

 

 

 

Rooi Hundu
0 q
. Dolomite o o o
0
° Limestone o 00
5 - °
1% 0
° 0
3‘
0 0
1O " O O
Depniln O» .
Meters . 9
15 q. . .
O 0
° 0
° 0
. O
20 «P .
O
O
O
O
25 c : as t : :

 

 

10 15 20 25 30 35 40 45 50
Porosity

Figure 31. Percent porosity versus depth for the carbonates from R001
Hundu. The limestone porosities have been corrected for calcite
cementation.

108

TABLE 22. POROSITY AND PERMEABILITY OF THE ROOI HUNDU CARBONATES (ARUBA,

 

 

NA)
SAMPLE DEPTH IN LITHOLOGY PERMEABILITY POROSITY METHOD
METERS (MlLLlDARClES)
A07A 1.0 LIMESTONE 35.4 PLUG.TS
A073 1.5 LIMESTONE 33.3 PLUG,TS
A070 2.0 LIMESTONE 33.1 PLUG,TS
5 2.5 LIMESTONE 4514.0 35.1 PLUG,TS
344 3.0 LIMESTONE 42.4 TS
A033 3.5 LIMESTONE 43.2 PLUG,TS
34-7 4.0 LIMESTONE 39.2 TS
14 4.5 LIMESTONE 4786.0 35.2 PLUG,TS
A09A 5.0 LIMESTONE 33.9 PLUG,TS
A093 5.5 LIMESTONE 39.1 PLUG.TS
A090 6.0 LIMESTONE 45.0 PLUG,TS
16 6.5 LIMESTONE 4634.0 37.3 PLUG.TS
A10A 7.0 LIMESTONE 42.7 PLUG,TS
A103 7.5 LIMESTONE 38.6 PLUG.TS
13 3.0 LIMESTONE 4273.0 33.9 PLUG.TS
34-10 3.5 LIMESTONE 33.9 TS
B4-13A 9.0 LIMESTONE 5.9 41.5 PLUG.TS
34-133 9.5 LIMESTONE 33.1 TS
34-14A 10.0 LIMESTONE 46.4 TS
20 10.5 LIMESTONE 6217.0 39.2 PLUG.TS
AO1A 12.0 DOLOMITE 26.7 Ts
A013 12.5 DOLOMITE 31.4 Ts
A010 13.0 DOLOMITE 2496.0 33.5 PLUG
A010 14.0 DOLOMITE 33.3 TS
A02A 15.0 DOLOMITE 23.2 TS
A023 15.5 DOLOMITE 25.3 TS
A020 16 DOLOMITE 594.0 35.5 PLUG
A03A 16.5 DOLOMITE 24.6 TS
A033 17.0 DOLOMITE 20.9 TS

Table 22 (cont'd.)

A03C
A03D
A04A
A048
A04C
AOSA
A058
AOSA
A068
A060

17.5
18.0
18.5
19.0
19.5
20.0
21.0
22.0
23.0
24.0

DOLOMITE
DOLOMITE
DOLOMITE
DOLOMITE
DOLOMITE
DOLOMITE
DOLOMITE
DOLOMITE
DOLOMITE
DOLOM ITE

109

19997.0 34.2
20.7
27.6
32.1
24.0
22.1
26.9
34.8
23.5
33.5

PLUG
TS
TS
TS
TS
TS
T8
T8
TS
TS

 

Limestone porosity average . 39.9%

Standard deviation - 3.9%

Dolomite porosity average - 28.6 %

Standard deviation - 5.4%

Permeability average - 4072md

Standard deviation - 2107md

Permeability average - 7696rnd
Standard deviation - 10.696md

T- test of the limestone porosity versus dolomite porosity

t-3.75
"-39

110

Origin of the Sero Domi Dolomites

Determining the origin of the Sero Domi dolomites will afford some
constraints on the type of dolomitizing solution to use for the bicarbonate
mass flux modelling. The origin of the dolomite in the Sero Domi
Formation on the neighboring island of Bonaire (Netherlands Antilles) was
suggested to have been due to either a refluxing brine (Def feyes et al,
1965; Bandoian and Murray, 1977) or an evaporatively concentrated fresh
water/seawater mix (Sibley, 1980). It is also possible, however, that the
dolomitizing solution was cold seawater (e.g. Saller, 1984a and Aharons et
al, 1987). The evidence presented is consistent with dolomitization by
either cold seawater or evaporatively concentrated brines.

The carbon isotopic values of the Sero Domi dolomites are lighter
than dolomites precipitated from seawater (Veiser, 1983 and Saller,
1984a). The light carbon signature of the dolomites could be the result of ;
1) carbon inherited from the limestones, 2) elevated temperature, and/or
3) the influence of organic carbon (soil gas or sulfate reduction) in the
dolomitizing solution. The negative carbon in the dolomites from 8001
Hundu (if it was inherited from the limestones) was probably due to the
influence of organic carbon from sulfate reduction. The amount of organic
carbon that needs to be incorporated into the limestones can be

determined from the following equation (Lawrence, 1973):
3'3 c of 6qu = mega-1193b (a)

where

111

111 = 111018 1131311011 01 0111111981110 carbonate

an : 3‘3 C of the carbon going into the authigenic carbonate

(1-111) = mole fraction of original carbonate

ab = 6‘3 C of original carbonate

Curtis (1978), and Presley and Kaplan (1968) reported that organically

08111180 00111011 from 1.118 81111018 FOOUCUOH 20118 1811983 110111 ’21 1.0 ’25 0’00

P08. With this range of values, approximately 15 percent of the carbon

present must be authigenic in nature to produce a limestone with a carbon

isotopic composition of - 3 °/oo PDB. The negative carbon in the dolomites

could therefore, be explained by the carbon in the precursor limestone.

Alternatively, the negative carbon values are probably not the result
of an elevated temperature, because this is not consistent with the heavy
oxygen isotopes observed.

The negative carbon could be from the dolomitizing solution. The

light carbon could be due to the influence of soil gas Co2 from a fresh

water source, or to light carbon evolved during sulfate reduction. The
heavy oxygen isotopes in the dolomites are not consistent with an
influence of fresh water. The light carbon and the heavy oxygen isotopes
could be due to sulfate reduction occurring in cold marine waters. The
amount of authigenic carbonate that would need to be precipitated (see
above) is not unreasonable considerl ng the amount of cement that was
precipitated at Root Hundu.

The oxygen isotopic values of the Sero Domi Formation dolomites are
enriched in 3‘8 o as compared to dolomites precipitated from normal sea

112

water. This enrichment in oxygen isotopic values could be due to the .
influence of either dolomitization by evaporatively concentrated waters,

or cold sea water. The temperature of the dolomitizing solution (with an

initial composition of 0 0/0031‘101'11 would be about 16° C (equation from

Matthews and Katz, 1977) to produce a dolomite with an oxygen isotopic

composition of +3.5 °/00 P08.

At least two models can be proposed for the occurrence of positive
oxygen isotopes and negative carbon isotopes in these dolomites. First, an
isotopically heavy water passed through an isotopically light substrate.
Oxygen will exchange 3 to 1 versus the carbon in the carbonate structure
and would therefore equilibrate with the solution's isotopic signature
more rapidly than carbon. In the second model, dolomitization is due to
cold seawater in the sulfate reduction zone. The clustering of oxygen
isotopes suggests a fluid with a relatively constant isotopic signature
over a few hundred thousand years (the time necessary to precipitate the
amount of cement observed at 8001 Hundu by sea water - see the
discussion of flux modelling below). A constant isotopic signature for
this prolonged time period without a large constant reservoir like the
ocean is not likely. Sabkha hydrology demonstrates widely varying oxygen
isotopic signatures from their landward extent to their seaward extent.

The trace element composition of dolomites has often been used as a
tool to determine the origin of the dolomitizing solution. The usefulness
of trace elements however is hampered by an inability to accurately
determine distribution coefficients for species in dolomites at low
temperatures. As a first approximation, strontium data were collected in

order to model the origin of the dolomites by looking for consistent trends

113

in the data. A compilation of isotopic and strontium compositions in some
relatively recent dolomites is presented in table 23.

The distribution coefficient of strontium in dolomite is still debated
(Land, 1980). Katz and Matthews (1977) suggest a distribution coefficient
of 0.025 from experimental evidence, whereas Baker and Burns (1985)
determined a distribution coefficient of 0.06 from Recent sediments.
Using these values as upper and lower limits, a dolomite precipitated from
seawater should have a strontium concentration of between 120 and 280
ppm. The concentration of strontium in the Sero Domi dolomites (176-259
ppm) f alls within this range and suggests the possibility of seawater as a
dolomitizing solution. These values are also similar to the concentrations
of strontium in the dolomites from the Eniwetok and Niue Atolls (1' able
23) which are both thought to have been dolomitized by seawater.

Refluxing brines should have higher concentrations of strontium due
to the evaporative concentration of the strontium in the waters. Dolomites
formed from this water should therefore have higher concentrations of
strontium (assuming a constant distribution coefficient). Dolomites
theorized to have formed in this way however, do not show any significant
difference in strontium concentrations from dolomites theorized to have
formed from seawater (T able 23). Therefore, strontium cannot be used as
a discriminator between the two diagenetic environments.

The similarity of the iron and zinc concentrations in the dolomite
also supports the premise that there was not a major difference in
solution chemistry.

The evidence presented is consistent with dolomitization by either
cold seawater 0r evaporatively concentrated brines. A dolomitizing

solution of seawater composition is however, more consistent with the

114

TABLE 23. STABLE 150T OPIC (FOB) Aw TRACE ELEI‘ENT CCI'POSITIONS(PPH) 0F
SELECTED CENOZOIC DOLMITES

 

 

Location 5r re Zn 3'30 3'90 ProposedOrigin
Jamalce-Falmouth 3000 -8.4 -1.0 Mixed
(Land, 19730)

Jamalce-l-lopegate 220 +12 +2.2 Mixed
(Land,1973a)

Bonalre 210 +2.4 +3.1 Mixed
(Sibley,19606Land, 1973a)

Colette Atoll 170 +40 +40 Mixed
(Rapellln,l977)

El Natl) 96 -0.5 -l.1 Mixed

(Land 31 al, 1975)

 

 

Eniwetok Atoll 170 36 +23 +2.5 Cold Marine
(Saller,1984a)

Niue Atoll 223 1587 151 +2.0 +2.8 Marine
(Aharon et el, 1987)

Sugarioaf It , FL 620 -0.7 +3.0 Marine
(Behrene and and, 19726.

Carballoet al, 1987)

Belize 600 - 1.6 + 1.6 Marine
(Mazzullo et al, 1987)

Jamaica 600 + 1.6 +2.0 Marine!
(Mitchell et al, 1987)

Montages Fm, CA 236 -6.13 +0.73 Marine
(Burns Baker, 1987) 558 +8.62 +0.50 Marine
Bahama Peri-platform 850 +2.82 +5.18 Cold Marine
(Mullins et al, 1985) . (Authegenic)
Israel Cretaceous 97 +0.49 -0.33 Marine-
(Sassandketz, 1982) Hwersallne
Sabkha 1 15 +0.4 swkha

(T 1196 M’Rabet, 1985)

Holocene Baffin Bay, 1'! 916 -0.9 +48 Mersellne
(Behrene and Land, 972)

Persian Bull 640 +4.3 +3.7 Evaporatlve

(Behrens and Land, 1972) Brine

115

clustering of oxygen isotopic data.

Discussion of the Bicarbonate Flux Model

The Sero Domi Formation on the island of Aruba was probably
dolomitized by either cold sea water or a refluxing brine. In order to
evaluate the mass flux models, a time constraint on the time allowable for
dolomitization needs to be determined. The Sero Domi dolomites are
Pliocene in age and the overlying Pleistocene limestones were not
dolomitized. The dolomitization then occurred in < 3myr. Mass flux
modelling is used to estimate the amount of time necessary to precipitate
the amount of dolomite cement observed in the Sero Domi by both a
refluxing brine and normal seawater. The average porosity of the
limestones at F1001 Hundu is 39.9 percent (corrected for calcite cement)
and the section is approximately 15 meters thick. If the limestones were
replaced by dolomite with a 13 percent volume loss of the solids, then
19.6 percent dolomite cement must have been added to account for the

present dolomite porosity of 28.1 percent. If sea water was the
dolomitizing solution, and HCD3‘ the controlling anion for cementation,
then the required amount of dolomite cement could have been precipitated
in .l -.34 million years, f luxing HCO3‘ by Kohout convection. This amount
of dolomite cement could be precipitated in .03-2 million years f luxing
HCO3‘ by sea water circulation associated with a mixing zone. if a brine
was the dolomitizing solution, the required amount of cement could have

been precipitated in .11 - 1.14 million years fluxing HCD3" by a refluxing

116

brine. As compared to the time available for dolomitization, the short
time periods obtained from the modelling, support the idea that the flux of
bicarbonate through the system is not the limiting factor to be considered

in the porosity evolution of these dolomites.

Eocene Carbonates From The Eniwetok Atoll

The Eocene strata from Eniwetok Atoll (Fig. 32) are lime mudstones
to grainstones with a thin horizon (B m) of dolomite. The dolomite is
planar and uniform with rhombs ranging from 50 to 310 microns in
diameter, and averaging of 137 microns. Partially dolomitized rocks
(coralline algal, f oram grainstones and packstones) contain
f abric-selective replacement dolomite and dolomite cement. The fabric
selective dolomite preferentially replaces the coralline algae and mud
(Figure 33). Dolomite cement is superimposed on calcite cement (Figure
34) and compaction fractures (Saller, 1984a). The dolomite in the
partially dolomitized rocks is the same size as the dolomite in the
completely dolomitized rock. The stable isotopes of the dolomite are 2.5
18o and 2.3 I3C (Saller, 1984b). The dolomites are Ca-rich with 55.3 mole

percent CaC03 and average 170 ppm strontium (Saller, 1984b). The

porosity of the limestones ranges from 33.2 to 9.8 percent and averages
23.1 percent (Figure 35 and Table 24). The permeabilities of the
limestones range from 2042 to 1658md (T able 24). The porosity of the
dolomitized rocks ranges from 14.2 to 4.2 percent and averages 7.4
percent. Only one permeability measurement was taken on the dolomites

and it was 1658md. The porosity of the partially dolomitized rocks

117

 

.:3< 3.2655 2: EB. no»: 88 8.: 02 SE 5:83 am 959...

 

 

 

illllfllllll‘
asapuIOJ; \

 

I I

d
a p as c

z
. z p-.. .30;

:24 5.2-.6 ..v

 

 

 

 

 

118

 

Figure 33. Selective replacement of the mud by dolomite and the
resistance of the coarsely crystalline fossil fragments. Width of field :
3.8mm.

119

 

Figure 34. Dolomite cementation (A) after calcite cementation (8) on
Eniwetok. Width of field : 0.2mm.

120

Eniwetok

12601'
128000,, . . . . «.1»

1300 "

 

m I o o e O Limestone

, 1320 1’ woo" '

Depth In 0 10-50% Dolomite
. I >50% Dolomite

7

Motors 1340 ‘

 

 

 

1360 1

“I

O

13301 °

 

1400 i 3 F
O 5 10 15 20 25 30 35
Porosity

1p
4
1

Figure 35. Percent porosity versus depth of the carbonates from the .

Eniwetok Atoll. Note the general decrease in porosity with the increase in
percent dolomite.

121

TABLE 24. POROSITY AND PERMEABILITY OF THE ENIWETOK ATOLL CARBONATES

 

 

SAMPLE DEPTH lN LITHOLOGY PERM EABILITY POROSITY METHOD
METERS (MILLIDARCIES)

F1 -1 1 -4 1279.0 LIMESTONE 1921.0 33.2 PLUG
F1 -1 1-5 1280.0 LIMESTONE 16.0 TS

F1 -1 1 -8 1281.0 LIMESTONE 1928.0 33.6 PLUG
F1 -1 1-9 1281.4 LIMESTONE 13.4 TS

F1 -1 1-12 1282.0 LIMESTONE 2042.0 32.2 PLUG
F1 -1 1-16 1282.3 LIMESTONE 1675.0 30.8 PLUG
F1 -1 1-18 1282.6 LIMESTONE 9.8 TS
F1 -1 1-21 1283.2 LIMESTONE 1686.0 29.1 PLUG
F1 +1 1-25 1283.5 LIMESTONE 1695.0 27.7 PLUG
F1 -1 1-26 1283.8 10-50 03 14.1 TS
F1 -1 1-30 1285.0 10-50 08 9.7 TS

F1 -1 1-34 1286.3 LIM ESTONE 22.4 TS

F1 -1 1-36 1286.9 LIM ESTONE 1501 .0 32.8 PLUG
F1 -12-1 1315.5 LIMESTONE 17.8 TS

F1 -12-2 1315.6 LIMESTONE 20.3 TS

F1 -12-3 1315.8 LIMESTONE 12.4 TS

F1 ~12-4 1316.1 >50 08 1428.0 8.7 PLUG
F1 -12+4T 1316.1 >50 OS 10.2 TS

F1 -1 2.48 1316.3 >50 OS 7.6 TS

F1 ~12-5T 1316.4 >50 03 5.2 TS

F1 -12-58 1316.4 >50 OS 7.4 TS

F1 -12-6 1316.6 >50 DS 1005.0 7.1 PLUG
F1 ~12-7T 1316.7 >50 08 2.8 TS

F1 -12-78 1316.9 >50 OS 6.1 TS

F1 -12-8 1317.3 >50 OS 4.2 TS

F1 -12-9 1317.5 10-50 08 1658.0 29.1 PLUG
F1-12-10 1317.7 >50 08 14.2 TS
F1-12-12 1318.7 10—50 OS 10.4 TS

F1 -12-14 1319.2 10-50 08 7.7 TS

F1 -12-16 1320.9 10-50 OS 11.6 TS

Table 24 (cont'd.)

F1-12-23
F1 -13-2
F1 .142
F1 -14-4
F1-14-8
F1-14-14
F1 -14-17
F1 4420
F1 -14-25
F1-14-28
F1-14-29
F1-15-3
F1-15-13

1322.8
1345.7
1371.6
1372.5
1373.7
1374.0
1374.7
1375.0
1376.5
1377.4
1377.7
1380.1
1383.0

122

1050 OS

LIMESTONE
LIMESTONE
LIMESTONE
LIMESTONE
LIMESTONE
LIMESTONE
LIMESTONE
LIMESTONE
LIMESTONE
LIMESTONE
LIMESTONE
LIMESTONE

1413.0
1106.0
1699.0

1634.0
1713.0

8.1
22.5
10.5
30.2
29.0
29.5
23.2
23.7
21.5
13.4
25.6
16.1
24.0

TS
T8
T5
PLUG
PLUG
PLUG
TS
PLUG
PLUG
TS
T8
T8
T8

 

Limestone porosity average - 23.1%

Standard deviation .. 7.5%

10-50% Dolomite porosity average - 13.0%

Standard deviation - 7.4%

>50% Dolomite porosity average - 7.4%

Standard deviation - 3.2%

Permeability average - 1668md

Standard deviation - 250md

Permeability - 1658md

Permeability average - 1216md

Standard deviation - 2991110

123

appears to vary inversely with the amount of dolomite (Figure 35 and
Table 24).

The fabric selective nature of the partial dolomites and the
similarity of crystal size of all the dolomites is consistent with the
hypothesis that the amount of dolomite in the rocks may have been
controlled by the original rock texture. As in many ancient rocks, the
dolomite may have selectively replaced mudstones. The fact that there
are only rare fossil ghosts and molds in the completely dolomitized
section is consistent with this hypothesis. If the original texture of the
dolomites and limestones were different, then one can only speculate as to
the original porosity of the sediments that were dolomitized. One can
estimate the predolomitization porosity from the current porosities in the
nondolomitized portion of the section. Enos and Sawatsky (1981). through
their work on Holocene sediments, show that sediments containing higher
proportions of mud were at least as porous as, or more porous than,
sediments containing higher proportions of allochems. If this is the case,
and only minor compaction and cementation is observed in the limestones,
it is reasonable to assume that the low porosity dolomite that replaced
the mud has resulted in a porosity loss of approximately 20 percent over

the precursor limestone.

Discussion of the Bicarbonate Flux Model

TO 1110081 1118 111111 01 H003” 101' 111888 FOCKS 1'801111’88 08811100110118

about the porosity of the rocks that were dolomitized. The present

porosity of the dolomites is 7.4 percent, whereas the the present porosity

124

of the limestones is approximately 24 percent. We will assume the
predolomitization porosity of the limestones was approximately 24
percent. 11 the rocks were replaced with a 13 percent volume reduction,
then approximately 26 percent dolomite cement must have been
precipitated. Saller (1984a) presented strong evidence that the
dolomitizing solution was sea water and that the dolomitization occurred

within the past 24 myr. lf sea water was the dolomitizing solution, and

H1303: 1118 COMFOIHTIQ 0111011 101' 08018111011011, 111811 1118 1'801111'80 0111011111 01

dolomite cement could have been precipitated in .1 to .24 million years

f luxing sea water by Kohout convection. This amount of cement could have

been precipitated in .02 to .14 million years f luxing HCO3' by sea water

circulation associated with a mixing zone.

For most dolomites, the flux models are difficult to constrain. In the
dolomites from Eniwetok, the dolomitizing solution and the maximum time
allowable for dolomitization are relatively well constrained. In this case,
we estimated a considerable porosity reduction within a time limit of .2
million years. As compared to the time available for dolomitization, the
short time periods obtained from the modelling again support the idea that
the flux of bicarbonate through the system is not the limiting factor to be

considered in the porosity evolution of these dolomites.

DISCUSSION OF DOLOMITE POROSITY

The ability to determine which horizons within a dolomite are going

to have a higher porosity is a very important in both reservoir exploration

125

and its subsequent development.

Early work on dolomitization and porosity concentrated on why
dolomites are more porous than limestones (Van Tuyl, 1914;11urray, 1930;
Landes, 1946; l'lurray, 1960 and Weyl, 1960). These models were based on
the assumption that the dolomitizing solution contained little or no
bicarbonate available for dolomite cementation. The flux modelling for
the Eniwetok and Aruba dolomites suggests that the flux of bicarbonate is
not always a limiting factor in the porosity evolution of a dolomite. The
condition of porosity development with dolomitization is therefore not
always the case. Often, coeval dolomites have a similar porosity to, or
lower porosity than, their coeval dolomites (Schmoker and Halley, 1982 ;
Halley and Schmoker, 1983 ; Barrett, 1987 and this study).

The Trenton Formation dolomites have a higher porosity than their
associated limestones. The dolomites however, are not 13 percent more
porous than the limestones, which is what would be predicted by the ‘local
source theory“, but are only approximately two percent higher.

A fundamental problem with most studies of dolomite and porosity is
that bulk dolomite and limestone porosities are compared. This causes
problems in that porosity in limestones and in dolomites can vary on the
scale of a thin section as well as between regional differences in f acies.
Dolomite porosities in the Trenton varied widely in the space of a few
centimeters. Any attempt to truely assess the role that dolomitization
plays in the evolution of porosity between limestones and dolomites needs
to compare similar initial f acies, that have been dolomitized by solutions
with similar chemistries. For example, Schmoker et al (1985) compares
dolomites (the bulk of which are from mid-lower Paleozoic) with

limestones that are generally younger. ‘With this type of analysis, one can

126

not assess the differences in diagenetic imprints over time. Also it is not
possible to assess porosity variances due to differences in original facies
or due to differing dolomitizing solutions.

Perhaps most dolomites initially have porosities similar to or even
lower than their coeval limestones. Subsequent to the dolomitization
episode, the unreplaced limestones will lose their porosity at a more rapid
rate due to later cementation, or differences in pressure solution.
Schmoker and Halley (1982) show that limestones from south Florida lose
their porosity faster than their coeval dolomites due to the 'porosity

reducing effects of burial".

SUMMARY

Dolomite porosity is the result of the interplay of factors influencing
reaction kinetics. These factors include: I) the surface area of the
substrate, 2) the saturation state of the solution, 3) temperature, 4)
substrate mineralogy, and 5) the effect of inhibiting ions. The flux of '
bicarbonate also affects dolomite porosity in that it influences the
amount of dolomite that can be precipitated.

The relationship between the bicarbonate flux, the surface area of the
substrate, the solution chemistry, and the other factors influencing
reaction kinetics and dolomite cementation are not clearly understood.
The precise conditions under which any of the other kinetic effects listed
may become the dominant textural (and therefore porosity) control are not
known.

The affects that the substrate surface area, the saturation state of

127

the solution, and the flux of solute had on the resultant porosity of a
dolomite were studied. The flux of bicarbonate determines the amount of
dolomite cement that may be precipitated. The surface area of the
substrate affects the porosity of a dolomite, in that it influences the
nucleation rate of the dolomite. A high surface area substrate leads to a
high nucleation rate and a low surface area leads to a low nucleation rate.
A high nucleation rate leads to tightly packed dolomite crystals. A low
nucleation rate leads to more loosely packed dolomite crystals. The
saturation state of the solution also has an influence on the nucleation
rate of dolomite. A high saturation state leads to a high nucleation rate,
and a low saturation state leads to a low nucleation rate.

In the Trenton dolomites and the Sero Domi dolomites, the surface
area of the substrate has a control on the resultant porosity of the
dolomite. The dolomitization of a finely crystalline (high surface area)
limestone resulted in dense dolomite nucleation and a dolomite with
porosities similar to or lower than the host. Dolomitization of a more
coarsely crystalline (low surface area) limestone resulted in less dense
dolomite nucleation and a dolomite with porosities higher than the host.

In the Trenton Formation dolomites, a change in fluid chemistry of the
dolomitizing solution had no significant influence on the resultant
porosity of the dolomites. This is not to say, however, that solution
chemistry changes would never have an influence on the dolomite porosity.
Perhaps the change in saturation state that the Trenton Formation was
subjected to was simply not large enough to produce any significant
difference in porosity.

The flux of the solute was not a limiting factor in the porosity

evolution of the dolomites from the Eniwetok Atoll or the Sero Domi

128

Formation. The amount of dolomite cement present in these rocks could
have been precipitated in geologically reasonable time frames by

reasonable concentrations of bicarbonate in solution.

CONCLUSIONS

The porosity of a dolomite is the result of the interplay of many
factors that influence reaction kinetics. These factors control the texture
and therefore the porosity of the resultant dolomite. The flux of
bicarbonate through the system also affects dolomite porosity in that it
influences the amount of dolomite cement that can be precipitated. In this
study, the effect of two of factors influencing reaction kinetics; l) the
surface area of the substrate and 2) the saturation state of the solution
were examined. The first conclusion is that the surface area of the
substrate is a dominant control on dolomite texture and therefore porosity
in these rocks. Finer-grained substrates that were dolomitized, resulted
in dolomites with a low porosity. This occurred due to the high
concentration of dolomite nuclei and the subsequent impingement of the
numerous crystals with growth. Coarsely grained substrates resulted in
textural reorganization and the development of more porous dolomites.
This occurred due to the low concentration of dolomite nuclei, which led
to more coarsely crystal line dolomite rhombohedra. The surface area of
the substrate can be shown to affect other textural features in a dolomite.
These features include: i) rhombohedra size variations within a sample, 2)
fossil moldic porosity and 3) textural preservation and/or textural

obliteration.

129

Second, that the change in saturation state did not appear to have an
influence on the porosity of the dolomite.

Third, that the flux of bicarbonate through the system is not always a
limiting factor in the evolution of dolomite porosity.

The dolomitization process can either create, mimic or destroy
porosity. The ultimate porosity of the dolomite is the result of the
interplay between differences in the substrate surface area and the

chemistry of the dolomitizing solution.

APPENDIX I

APPENDIX I

COI‘IPARISON 0F COHPENSATED NEUTRON LOG POROSITY
WITH WHOLE CORE POROSITY

The following is a comparison of the whole core porosity with the
compensated neutron log porosity. All of the logs in the area were run by
Schlumberge, so that no correction is needed in comparing logs from
different wells. In this comparison, since there are different correction
factors involved in determining the porosity of limestones and dolomites
from the logs, the simplest and least confusing way to compare the
porosities is to compare the neutron porosities with a core that has all
one lithology. In this case, the Total Petroleum Luck l-12 well was
chosen because the cored interval was completely dolomitized. In
comparing the porosities listed below, one must keep in mind that the
compensated neutron log porosities are an average porosity of the
formation near the borehole, whereas the whole core porosity is an
average over a smaller area. Therefore any localized deviations from the

norm will be more apparent in the whole core porosity values.

130

131

TABLE AI-l. COHPARISON OF COHPENSATED NEUTRON AND WHOLE CORE
POROSITY OF THE TRENTON DOLOMITES.

 

 

Depth Compensated Neutron Whole core
(F 900 Percent Porosity Percent Porosity
4842 0.9
4883 0.8
4844 0.5 1.1
4845 1.4
4846 0.9 1.0
4847 1.9
4848 1.7 2.2
4849 1.4
4850 1.7 2.0
4851 1.9
4852 2.0 2.4
4853 2.8
4854 2.5 2.7
4855 2.6
4856 3.0 2.9
4857 3.5
4858 2.5 5.3
4859 4.5
4860 2.0 6.4
4861 3.7
4862 4.4 4.3

4863 2.7

Table Al-l (cont'd.)

4864
4865
4866
4867
4868
4869
4870
487 1
4872
4873
4874
4875
4876
4877
4878
4879
4880
488 1
4882
4883
4884
4885
4886
4887
4888

132

10.1

7.5

4.4

6.8

3.9

3.9

3.0

3.0

3.0

3.4

2.5

5.0

2.5

2.9
1 .8
1 .6
8.9
5.9
8.9
3.8
3. 1
3.3
5.2
8.3
6.4
2.9
3.2
3.4
3.7
3.6
3.5
3.9
2.7
3.7
4.0
3.3
4.8
3.3

Table Al-l (cont'd.)

4889
4890
4891
4892
4893
4894
4895
4896
4897
4898
4899
4900
4901
4902
4903
4904
4905
4906
4907
4908
4909
4910
4911
4912
4913

133

31

122

L4

L4

21)

21)

21)

51)

31)

122

21)

21)

I36
I32
I33
2A1
‘43
I36
I34
I36
40
25
27
29
4A
35
38
23
30
37
48
30
23
32
32
38
31

Table Al-l (cont'd.)

4914
4915
4916
4917
4918
4919
4920
4921
4922
4923
4924
4925
4926
4927
4928
4929
4930
4931
4932
4933
4934

£9.25

Mean
51)

134

213

:25

31)

51)

‘44

31)

21)

:25

:25

:25

‘44

31
L8

Mean
511

31)
27’
31)
134
511
71
3A1
31)
31
‘46
5A1
51
138
I36
:25
31)
21
21
:23
215
215
§J_
235
L5
93

135

A student T test indicates that the two methods of porosity analysis are

not significantly different.

APPENDIX 11

APPENDIX 11

WHOLE CORE POROSITY AND PERMEABILITY ANALYSES OF THE
TRENTON FORMATION CARBONATES

TABLE All - 1. WHOLE CORE DOROSITY AND DERMEABILITY ANALYSES OF THE TRENTON
FORMATION.

 

5 10 th Porosit Permeabilit Lithol
m Dep V Max. 90 Deg. 09y

 

Total Luck 1-12

1 40430-440 0.0 0.2 <01 Ds,Sl/Anhy
2 40440-450 1.1 <01 <01 Ds,Sl/Anhy
3 40450-400 1.4 <01 <01 Ds,Sl/Anhy
4 40400-470 1.0 <01 <01 03,51/Anhy
5 4047.0-400 1.9 0.0 0.2 Ds,Sl/Anhy
0 4040.0-490 2.2 0.2 0.1 08,5l/Anhy
7 40490-500 1.4 0.4 <01 03,51/Anhy
0 40500510 2.0 <01 <01 03,51/Anhy
9 40510-520 1.9 <01 <01 0351771111117
10 40520-530 2.4 <01 <01 0051711111117
1 1 40530-540 2.0 <01 <01 Ds,SI/Anhy '
12 40540-550 2.7 0.1 <01 03,51/Anhy
13 49550-500 2.0 0.1 <01 03,51/Anhy
14 4050.0-570 2.9 0.1 <01 Ds,Sl/Anhy
15 40570-500 3.5 <01 <01 03

10 40500-590 5.3 1.0 0.4 03

17 40590-000 4.5 0.5 0.1 03

10 40000-010 0.4 1.0 1.0 03

19 40010-020 3.7 0.3 0.1 03

20 40020-030 4.3 0.7 0.5 03

21 40030-040 2.7 9.2 1.5 05

22 40040-050 2.9 1 1.0 0.1 05

23 40050-000 1.0 <01 <01 0:

24 40000-070 1.0 * <01 03

25 40070-000 0.9 27.0 9.1 DsPyrite

136

Table AII-l (cont'd)

26

1'0
‘1

28333385882833888

83$

83285338853153?!

48680-690
48690-700
4870.0-71 .0
4871 0-720
48720-730
48730-740
48740-750
48750-760
48760-770
48770-780
4878.0-790
48790-800
4880.0-81 .0
4881 0-820
48820-830
48830-840
48840-850
48850-860
48860-870
48870-880
48880-890
48890-900
48900-91 .0
4891 .0-920
48920-930
48930-940
48940-950
48950-960
48960-970
48970-980
48980-990
48990-000
49000-01 .0
4901 .0-020
49020-030
49030-040
49040-050
49050-060

5.9
6.9
3.8
3. I
3.3
5.2
8.3
6.0
2.9
3.2
3.4
3.7
3.6
3.5
3.9
2.7
3.7
4.0
3.3
4.8
3.3
3.6
3.2
3.3
2.4
4.3
3.6
3.4
3.0
4. l
2.5
2.7
2.9
4. I
3.5
3.8
2.3
3.0

137

1 0.0
6.6
0.4

<0. 1

<0. I
0. 1
0. 1

1 5.0

<0. 1

<0. 1
0.5
2.0

<0. I
0.5

<0. I

<0. I

<0. I
0.3
0.3

63.0

<0. I

<0. 1
<0. I
0. I
<0. 1
0. 1
<0. I
0.6
<0. I
<0. I
<0. I
0.3
0.3
0.5
0. 1
<0. I

5.7
4.4
0.2
<0. I
<0. I
<0. I
<0. 1
9.7
<0. I
<0. I
0.4
<0. I
<0. I
<01
<0. I
<0. 1
<0. I
0.2
0.1
0.2
<0. I
<0. I
<0. 1
<0. I
<0. I
<0. 1
<0.1
<0. I
<01
0.4
<0. I
<0. I
<0.1
0.1
0. 1
0.4
<0. I
<0. I

Ds,1>yrite
Ds,l>yr'ite
Ds,l>yrite
Us

05

9989893939888

5
<
#1

S’EES’EE’ESSS’S’E’S’S’SES’S’

138

Table All-l (cont'd)

64 4906.0-070 3.7 0.1 <0. I 03

65 49070-080 4.8 2.8 1.1 08

66 49080-090 3.0 <0. I <0. I 03

67 49090-100 2.3 0. I <0. I 03
4910.0-1 I .0 3.2 1.9 <0. I Os,V/F

69 491 I .0- 12.0 3.3 0.2 0. I 03
49120-130 3.0 0. I <0. I 03

71 49130-140 3.1 <01 <01 03
4914.0- 15.0 3.0 <0. I <01 08
4915.0- 16 0 2.7 <01 <01 03

74 4916.0- 17.0 3.0 0. I <0. I 03
49 I 7.0- 1 8.0 3.4 0. I <0. I 03

76 4918.0- 19.0 5.4 2.7 <01 03
49 19.0-20.0 7. I 18.0 2.0 05,51/Lmy

78 4920.0-210 3.4 <0. I <0. I 08

79 4921 .0-220 3.0 0.5 0.2 OS

80 49220-230 3. I <0. I <0. I D8,SI/Lmy

81 49230-240 4.5 0.3 0.2 03

82 49240-250 5.4 0.5 0.3 Ds,SI/Lmy

83 4925.0-26 0 5. I 0.3 0. I 03

84 49260-270 3.8 0.5 0.3 08

85 49270-280 3.6 <0. I <0. I 03

86 49280-290 2.5 <0. I <0. 1 OS

87 49290-300 3.0 <0. I <0. I D3

88 4930.0-310 2.1 <0. I <0. I 03

89 4931 0-320 2.3 <0. I <01 D3

90 49320-330 2.6 0.2 0. I 03

91 49330-340 2.5 0.4 0. I 03

92 49340-350 3. I 0.4 0.3 03

Tgtgl Emil F ist -l2

1 50240—250 0.2 0.2 0.1 10517030175011)
2 50200-270 0.2 <01 <01 13,5170».

3 50200-290 0.4 0.2 <01 15517001771

4 50300-310 0.4 <01 <01 13,5170:

5 50310-320 0.2 <01 <01 Ls,SI/Ds

139

Table All-l (cont'd.)

6 50320-330 0.3 <0. I <0. I Ls,SI/Ds
7 50330-340 0.4 15.0 <0. I Ls,51/Ds,V/F
8 50340-350 0.3 <0.l <0. I Ls,Sl/Ds,V/F
9 5035.0- 36.0 0.5 <0. I <0. I LsSl/DsN/F
10 5036.0-370 0.4 <0. I <0. I Ls,Sl/Ds,V/F
l l 50370-380 0.2 <0. I <0. I Ls,SlIDs
l 2 50380-390 0.3 <0. I <0. I Ls,SIIDs
l 3 50390-400 0.2 <0. I <0. I Ls,Sl/Ds
1 4 5040.0-4 I .0 0.3 <0. I <0. I Ls,SI/Ds
15 504 1 .0-420 0.4 <0. I <0. I Ls,Sl/Ds
1 6 50420-430 0.2 <0. I <0. I Ls,SI/Ds
17 50430-440 l .2 5.4 <0. I Le,SI/Ds,V/F
l 8 50440-450 0.3 <0. I <0. I Ls,Sl/Ds
l 9 50450-460 0.7 ’1‘ <0. 1 1.551105
20 50460-470 0.4 <0. I <0. I LsSI/Ds
21 5047.0-480 0.2 ’“ <0. 1 L5
22 50480-490 0.4 <0.1 <0. I Ls,Sl/Ds
23 50490-50 0 0.7 <0. I <0.l Ls,Sl/Ds
24 5050.0-5 I .0 0.? <0. I <0. 1 L503
25 5051 0-520 0.8 <0. I <0. I Ds,Lmy
26 50520-530 0.4 ’1‘ <0. l 1.3
27 50530-540 l .8 0| <0. I 051/151!ng
28 50540-550 1.7 <01 <01 05,51!ng .
29 5055.0-560 l . I <O.l <0. 1 05,511ng
30 50560-570 l .7 <0. I <0.l 03
31 50570-580 2.6 * <0. l 03
32 50580-590 2.7 * <0. l 03
33 50590-600 3.7 * <0. I 133
34 5060.0-61 .0 l .9 0. l ‘ <0. 1 03
35 5061 0-620 2. 1 <0.1 <0. l 03
36 5062.0-630 2.7 <0. I <0. I 05,51/Lmy
37 50630-640 2.3 "‘ <0. I 03
38 50640-650 2.8 "‘ <0. I 03
39 50650-660 2.8 ’1‘ <0. l 03
40 50660-670 2.3 'I‘ <0. I 05
41 50670-680 l .6 <0.l <0. I 03
42 50680-690 2.7 <0. I <0. I m,ng,Sty,V/F
43 50690-700 2.0 ’1‘ <0. I 05

Table AII-I (cont'd.)

832858318$€£339838

2311213388

:13ch

50700-7 I .0
5071 .0-720
50720-730
50730-740
50740-750
50750-760
50760-770
50770-780
50780-790
50790-800
50800-8 I .0
508 1 .0-820
50820-830
50830-840
50840-850
50850-860
50850-860
50870-880
50880-890
50890-900
5090.0-91 .0
50920-930
50930-940
50940-950
50950-960
50960-970
50970-980
50980-990
50990-000
51 00.0-0 I .0
51 01 .0-020
51 02.0-03.0
51 03.0-04.0
51 04.0-05.0
51 05.0-06.0
51 06.0-07.0
51 07.0-08.0
51 08.0-09.0

2.7
I .4
3.8
2. I
2.8
2.3
3.3
2.2
2.5
2.8
2. I
3.5
4.6
4.7
5.0
4.2
I. I
2.8
2.0
1 .4
0.4
2.6
2. I
I .2
3. I
3.5
6. I
0.9
2.2
3.0
1 .8
I .2
2. I
2.6
2.0
1 . I
3.4
2.5

140

I“

0.2
<0. I
<0. 1

fl

2680
<0. I
<0. 1
<0. I
<0. I
<0. I
<0. I
<0. I

3.4

0.2

<01
<01

<0. I
<0. I
<0. I
<0. I
<0. I

0.4
<0. I

<0. I
<0. I
21.0
702.0

98
2. I
1 09.0

<0. I
<0. I
<0. I
<0. I
<0.l
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0.l
<0.l
<O.l
<0. I
<0. I

0.4
<0.l

0.2
<0. I
<0. I
<0. I
<0. I
<0. I
16.0
<0. I

0.2
<0. I

Ls,Ds,V/F
L503
L503
1.305
L303
L503
L303
Ds,l.my
Ds,Lmy
Ds,Lmy,V/F
Ds

Day/F
05,Lmy
05,51/Lmy
Ds,Lmy
Ls

1.5,03
05,51/Lmy
05,51/Lmy
Ds

Ds,Lmy
Ds,Lmy,1>yr,V/F
Day/F

Da,WI-'

De,V/F
Ds,SI/ng,V/F
03,51 Ing,V/F

Table All-I (cont'd.)

82
83
84
85
86
87
88
89
90
9 I
92
93

883883

100
101
102
103
104
105
106
107
108
109
110
III
112
113
114
115
116
117
118
119

51 09.0- I 0.0
51 10.0- 1 I .0
51 I 1.0- 12.0
51 12.0- 13.0
51 130- I 4.0
51 140- 15.0
51 15.0- 16.0
51 16.0- 17.0
51 170- 18.0
51 180- 190
51 19.0-20.0
51 20.0-21 .0
51 2 I .0-220
51 22.0-23.0
51 23.0-24.0
51240-250
51250-260
51260-270
51 27.0-28.0
51 280-29 0
51290-300
51300-3 I .0
51 31 .0-320
5132.0-330
51330-340
51 340-350
51350-360
51360-370
51370-380
51380-390
51390-400
5140.0-4 1 .0
514 1 .0-420
51420-430
51430-440
51440-450
5 I 45.0-46.0
5146.0-470

2.6
7. I
1 0. 1
8.3
6.3
4.0
3.2
3. I
2. I
1 .3
2.3
5.3
3.4
2.7
2.4
I .7
2.4
2.5
4. I
I .7
4.5
2.3
3. I
5.3
2. I
3.6
3.6
1 .5
3.6
0.9
0.5
0.3
0.6
0.9
I . I
I . I
1 .5
2.3

141

X

563.0
4.7

202.0
0.4
30.0

1 27.0

254.0
0. I
84.0
<0. 1
4.6
6.5
<0. I

0.2
37.0
0.2
0.2
0. I
0.2
0. I
0.3
<0. I
0.3
<0. I
0.3
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
2.5
<0. I
I .7
<0. I
0.2
44.0
<0. I
19.0
2.0
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I

Day/F
05,51!ng,V!F
03,Lmy,51!ng,V!F
05,Lmy

05,Lmy,51 Ing
03,51!ng

03,51 !ng,V!F
05,51 !ng,V/F
051,51!ng
03,51!ng,V/F
05,51N9y
Ds,I.my,51!ng,V!F
Da,l.my,51!V9y
05,Lmy,51 Ing
05,Lmy,Sl!ng
Ds,Lmy,Sl/ng
05,Lmy,5l!ng,
Ds,Lmy,SI!ng,V!F
Ds,Lmy,SI !ng,V!F
Da,l.my,51 Ngy
Ds,Lmy,Sl!ng,V!F
03,51!ng
Ds,Lmy,Sl!ng,V!I-'
05,51Ngy,V/F
Ds,Lmy,SI!ng,V!l-’
05,Lmy,SI/ng,V/F
Ds,l.my,51!ng,V!I-'
Ls,SI!Ds

L5,Ds

1.303

L305

L5

L5

L5

L3

L5

L3

L5

Table AII-I (cont'dl

I20
I21
122
I23
I24
125
126
127
128
I29
130
I31
132
133
134
135
I36
I37
138
I39
140
I41
142
143
144
145
146
147
I48
I49
150
151
152
153
154
155
156
157

S I 47.0-48.0
51 48.0-49.0
51 49.0-50.0
51 500-5 1 .0
515 1 .0-520
51 52.0-53.0
51 53.0-54.0
51540-550
51 55.0-56.0
51 56.0-57.0
51 57.0-58.0
51 58.0-59.0
51 59.0-60.0
51 600-61 .0
51 61 .0-620
51 620-630
51 63.0-64.0
51 64.0-65.0
51 65.0-66.0
51 66.0-67.0
51 67.0-68.0
51 68.0-69.0
51 69.0-70.0
51 70.0-71 .0
51 71 .0-720
51 72.0-73.0
51 73.0-74.0
51 74.0-75.0
51 75.0-76.0
51 76.0-77.0
51 77.0-78.0
51 78.0-79.0
51 790-80 0
51 80.0-81 .0
51 81 .0-820
51 82.0-83.0
51 83.0-84.0
51 840-850

I .3
0.2
0.8
0.5
0.8
0.8
0.7
2. I
I .6
2. I
6.3
4. I
4.3
7.3
5.5
4.4
0.2
2.0
6.7
4.5
2. I
2.5
4.7
3.8
3.4
3.4
4.8
2. I
3.0
2.3
0.4
0.9
0.3
0.3
0.4
0.3
0.3
0.5

142

<01
<0. I
0.2
<0. I
<01
4.1
0.1
0.2
14.0
207.0
0.1
1.3
<01
0.1

<0. I
6.3
0. I
0.3
0.2
3.5

1 47.0

43.0

<0. I

<0. I
<0. I

0.4
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I

<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I

I .5
<0. I
<0. I

2.5

0.3
<0.l

0.2
<0. I
<0. I
<0. I
<0. I

2.0
<0. I
<0. I
<0. I

2. l

0. l

1.8
<0. I
<0. I
<0. I
<0.l
<0. I
<0. I
<0. I
<0. I
<0. I
<0.l
<0. I
<0.1
<0.1

L3

L3
Ls,SI!Ds
Ls,SI!Ds
Ls,SI!Ds
L505
L303
03
05,51!Lmy
Ds
Day/F
05,V!F
Ds
Ds,Lmy
05,Lmy
03,Lmy
Ds,Lmy

03
03
D:
03
Us
05

Ds,V!I-'
Ds,SI!Lmy,V!F
05,51!Lmy,V/F
05,51!Lmy

05

03

03

1.503

1.505

1.503

Ls,SI!Ds
Ls,Sl!Ds

L305

1.305

1.303

Table All-l (cont'd.)

158
159
160
161
162
163
164
165
166
167

168
169

I70
171
172
I73
I74
175
I76
177

51 850- 86.0
51 86.0-87.0
51 91 .0-920
51 95.0-96.0
51 99.0-00.0
5231 .0- 32.0
52340-350
52390-400
52430-440
52440-450
52450-46 0
52460-470
5247.0-480
52480-490
52490-500
52500-51 .0
5251 .0-520
52520-530
52530-540
52540-550

Total flermon-Lyck l-l

04001-50110‘

GAun~oo

49650-660
49660-670
49670-680
49680-690
49690-700
49700-71 .0
4971 .0- 72.0
49720-730
49730-740
49740-750
49750-760
49760-770
49770-780
49780-790
49790-80 0

0.3
0.4
0.7
0.5
0.2
4.3
3.5
2.2
0.7
0.6
2.2
5. I
9.2
3.3
2.0
1 .7
6.2
0.9
1 .4
1 .7

0.9
1.1
1.1
1.1
1.0
1.1
1.3
2.0
3.2
1.4
1.3
1.5
1.3
1.3
0.9

143

<0. I
<0.l
<0. I
<01
<01
<0. I
0.3
O. I
(0.1
<0.1
<0. I
0.5
201.0
<01
<01
<0.l
0.3
<01
<01
<01

***************

<0. I
<01
<01
<0. I
<01
<01
<01
<0. I
<0. I
<01
<0.1
0.3
172.0
<01
<0. I
<01
<0. I
<0. I
<01
<01

<0. I
(0.1
<O.l
<0.l
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0. I
<0.l
<0. I

L505
L5

L5

L5

L5

L5

L5

L5
Ls,SI!Ds
Ds,Lmy
Ds,Lmy
L505
05,Lmy,V!F
L5

Ls

L5
L505
05,Lmy
L5
L505

L5
L5
L5
L5
L5
L5
L5
L5
L5
L5
L5
L5
L5
L5
L5

Table AII-I (cont'd.)

49800-81 .0
4981 .0-820
49820-830
49830-840
49840-850
49850-860
49860-870
49870-880
49880-890
49890-900
49900-91 .0
4991 .0-920
49920-930
49930-940
49940-950

Tgtgl Hgnkehd 1- I 7

dOOOxIOUI-AOUN‘

484 1 .0-420
48420-430
48430-440
48440-450
4880.0-81 .0
4881 .0-820
48820-830
48830-840
48840-850
48850-860
4886.0-870

ME 1'81 I"

I

2
3

4896.0-970
48970-980
4898.0-990

I .2
1 .9
3. I
0.9
2.2
2.4
3.2
I .3
2.1
2.2
0.8
0.7
0.7
0.7
0.7

0.6
0.4
0.4
0.4
0.4
0.6
0.5
0.7
0.7
0.5
0.4

3.0
3.2
1 .7

144

*******¥*******

<0. I
<0.I
<0.I
<0.1
<0.1
<0.1
<0.l
<01
<01
<0. I
<0.l

<0. I
<01
<01

<01
<0. I
<01
<01
<01
<01
<01
<01
<01
<01
<01
<01
<01
<01
<01

<0. I
(0.1
<0. I
<0. I
<0. I
<O.I
<0.1
<0.1
<0. I
<0. I
<0. I

<01
<01
<01

L5
L5
L5
Ls,SI!Ds
L5
L5
L5
L5
L5
L5
L5
L5
L5
L5
L5

L5,SI!Ds
L505
L505
L505
L5

L5

L5
L5,SI!Ds
L5

L5

L5

Da,V!51!Lmy
05,V!SI!Lmy
05,V!51!I.my

Table All-l (cont'd)

Total Luck 2-12

O-QOUI-AOIM‘

21

88159539888388288

48560-570
48570-580
50170- I 8.0
50 I 8.0- I 9.0
50 I 9.0-20.0
50200-21 .0
5021 .0-220
50220-230
50230-240
50240-250
50250-260
50260-270
50270-280
50280-290
5029.0- 30 0
50300-31 .0
5031 .0- 32.0
50320-330
50330-340
50340-350
5036.0- 37.0
50370-380
50380-390
50390-400
50400-41 .0
504 1 .0-420
50420-430
50430-440
50440-450
50450-460
50460-470
50470-480
50480-490
50490-500
50500-51 .0
50520-530

3.4
3.0
1 .4
2.5
1 .8
0.9
1 .I
I .2
7.2
8.4
2.7
4. I
3.9
5.4
6.8
4.7
4.0
I .7
2.2
2.0
2.5
4.0
2.7
I .8
2.4
3.2
I .2
2.2
I .3
1 .3
2.3
3.2
2.2
1 .8
I .3
I . I

145

<0.l
<0.l

0.1
<0.1

0.1
<0.1
<01
<0.1

0.5
<0. I
<0.I
<0.l
<0.l

<01
<01
<01
<01
<01
<01
<01
<0.1

0.8

0.2
<01
<01
<01
<01
29.0

0.1
<01
<01
<01
<01
<0.1

0.1
<01
<01
<01
<01
<01
<01
<0.1
<01
<01

0.2
<01
<01
<01
<01

05

05

05

D:

05

L5, 05
L505
05,Lmy
Da,Lmy,SI/ng,V!F
05,Lmy,51!ng,V/F
Ds
Ds,Lmy
D5
05,Lmy
05
05,V!F
Da,V!F
Ds
05,Lmy
05,V!F
05

05

05

05

05

05

05

03
L305
L505
05

05

05
05,Lmy
L505
L505

Table All-I (cont'd)

£32183

42
43
44
45

8588385592539833

281:!

50540-550
5056.0- 57 0
5058.0-59 0
5060.0-61 0
50620-630
5064.0-650
5066.0-670
50680-690
5070.0-71 .0
50750-760
5077.0- 78.0
5080.0-81 .0
50820-830
5083.0- 84.0
50840-850
50850-860
50860-870
50870-880
5088.0-89 0
50890-900
5090.0-91 .0
5091 .0-920
50920-930
50930-940
250940-950
5095.0-96 0
50960-970
5097.0-980
5098.0-990
5099.0-000
511110-01 .0
5101 .0-020
51020-030
51030-040
51040-050
51050-060
51060-070
5107.0-080

1 .3
0.8
I .0
1 .0
0.9
I .0
0.6
0.6
0.5
1 .2
0.7
0.9
I .2
0.8
1 .2
I .0
0.7
0.7
0.9
1 .0
0.6
0.9
1 .8
3.8
0.7
I .0
0.9
I .2
I .4
I .5
I .6
0.5
0.8
1 .7
0.8
I .5
1 .8
1 .6

146

<0. I
<01
<01
<01
<0.I
<0.1
<0.1
<0. I
(0.1
<0.1
<0.l
<O.I
<0. I
<0.l
<0.l
<0.1
<0.1
<0.1
<0. I
<01
<01
<01
<0.I

1.8
<0.1
23.0
<0.l
<01
<0.I
<0.1
(0.1
<01
<01
<0.1
<0.1

0.1

0.1

0.1

<01
<01
<01
<0.1
<0.1
(0.1
<01
<01
<01
<01
<01
<01
<01
<01
<01
<01
<0.l
<01
<01
<01
<0.I
<01
<0.1

1.7
401
40.1
(0.1
<01
<0.I
<O.I
40.1
<0.1
<0.1
40.1
<0.I

0.1
<0.I

0.1

L505
L505
L505
L5,Sl!05
Ls,SI!Ds
L5,SI!Ds
L5,SI/Ds
L503
Ls,SI/Ds
L5

L5, 05
L503
Ls,SI!Ds
L505
L5.SI!05
leos
L505
L505
L5
Ls,Sl/D5
L505
L505
L505
L505
L505
L5,Ds,V/F
L505
L505
L505
L505
L505
L5,SI/Ds
L505
L505
L505
133le
L505
L505

Table All- I (cont'd.)

147

 

75 51080-090 1.6 0.1 <0.1 L505
76 51090- 10.0 2.7 0.7 0.6 L5
77 51 10.0- I I .0 2.0 <0.1 <0.1 L505
78 51110-120 2.1 0.1 <0.l L505
79 51 13.0- 14.0 I .5 <0.1 <0.1 L503
80 51 15.0- 160 I .8 86.0 (0.1 L5,Ds,V!F
81 SI 17.0-18.0 1 .0 <01 <01 L505
82 51 19.0-20.0 I .0 <0.1 <0.l Ls,SI!Ds
83 5121 .0-22 0 I .l <0.l <0. I Ls,SI/Ds
84 51230-240 1.2 <0. I <0.l L505
85 51250-260 0.7 <01 <01 L5
86 51270-280 l .3 <0.l <0.l L505
87 SI 29.0-30 0 I .S <0.l <0.l L505
88 5131 .0-320 1.4 <0.1 <0.1 L505
* Indicates plug permeability
Total Average
Dolomite Porosity 3.1
Standard Deviation 1.4
Number of Samples 170
Total Average
Limestone Porosity 1.3
Standard Deviation 0.9
Number of Samples 62
Total Average
Partial Dolomite Porosity 1.9
Stanchrd Deviation 1.9
Number of Samples 169
Symbols Used:
Anhy - ariiydrite SI - slight Sty - styolites
ng - vuggy Ds - dolomite L5 - limestone

5th - shale

APPENDIX III

APPENDIX III

W-TEST TO DETERMINE THE NORMALITV OF THE TRACE ELEMENT
DATA IN THE TRENTON DOLOMITES

The following is the statistical analysis to determine if the
geochemical data has a normal or lognorrnal distribution. This was done in
order to define which test would be used to determine if the data sets were
statistically different.

The W-test is a method for testing whether a data set has been drawn
from a normal distribution (Gilbert, 1987). If the test is conducted on the
logarithms of the data, the hypothesis of a Iognormal distribution can also
be evaluated. This test is good for data sets of 50 samples or less. All of
the samples tested were from the Total Luck 1-12 and 2-12 wells, because

these wells have the greatest number of samples.

111-5

The first step is to determine the null hypothesis;

Ho = the population has a normal distribution

00:200-102

2) Arrange data from the smallest to the largest

3) Compute K

148

149

K = n/2 if n is even
K = n-1/2 if n is odd

4) 0010111 0081110181118 0' "' 0k from 0 51011311001 10018

5) Compute
W : 1/0120i1XIn-l,j] ' xlfllz

6) Reject H0 at 1118a significance level if 11 is loss than the quartile given
in the table.

 

RDNT I 11
The strontium values are from the Total Luck 1-12 and Total Luck 2-12

cores (n: 29).

w = 1!10,343 [0.4291(203-90) + 0.2900(202-102) + 0.2499(100-122) +
0.2150(101-123) + 010041102425) + 0.1010(101-125)+
0.1395(159-127) + 0.1 1921 157-129) + 0.1002(155- 1291+
0.0022(155-1301 + 0.0050(150-134) + 0.0403(149-135)+
0.0320(144-139) + 0.0159(141-1391l2

177:1!113,343*{131}2

W = 0.936

150
W at (952 CI) = 0.926

The calculated W is greater than Wa, we therefore accept the null

hypothesis that this population has been drawn from a normal population.

0005111551;
The manganese values are from the Total Luck l-l2 and 2-12 cores
(11:25).

111 = 1!1,391,300 {0.4450I2079-12131 + 0.300912070-12001 +
0.254312001-12071 + 0.214011050-12901 +
0.102211019-13051 + 0.153911791-1425] +
o.120311777-14091+ 0.104011700-15101 +
0.002311730-15571 + 0.001011714-15091 +
0.040311700- 10221 + 0.0200110051020112

W: 1!1,391,300*{110012

W = 0.967

W at (958 CI) is 0.918

The calculated W is greater than Wa, we therefore accept the null

hypothesis that this population has been drawn from a normal population.

IRN

The iron values are from the Total Luck 1-12 and Total Luck 2-12 cores

151

(11: 25).

w = 1!7.14 x 1051.4493130400-15070) + .3090134910—10221) +
.2554129329-17770) + 2145(20200-10100) +
.1007(27230- 10329) + .1512127143-10329) +
. 124512509 1 - 10079) + .0997125057- 10913) +
.0704124503-19090) + .0539123050-19202) +
0321(22077-19034) + 0107(221 12-21015) I2

w = 1!7.14 x 10312540012

1'1 : .904

W at (988 C1118 0.901

The calculated W is greater than Wa, we therefore accept the null

hypothesis that this population has been drawn from a normal population.

11.112

The zinc values are from the Total Luck 1-12 and 2-12 cores (n=25).

v1 = 1!123 1.4450(10-1) + 3009(9-2) + .254310-2) + .214010-2) +
.102215—2) + .153915-2) + .1203(5-2) + .104015-2) +
.002315-2) + .001015-2) + 0403(4-3) + 0200143)]?

14:1!12319)2

152
W: .660

W at (958 C1118 0.918

The calculated W is lower than Wa, we therefore reject the null

hypothesis that this population has been drawn from a normal population.

Taking the natural log of the values and pref orming the same test results

in:

w .-. 1/702 1.4450(23-00) + .3009122-09) + .25431179-09) +
.2140(1.79-.09) + .1022(1.01-.09) + 153911.01-09) +
.12031101-09) + .104011.01-.09)+ 002311.01-09) +
001011.01-09) + 040311.39-110) + .0200(I.39-1.10)l2

v1 = 1!7.02 [2.7112
W = .939
W at (952 CI) is 0.918
The calculated W is greater than Wu, we therefore accept the null

hypothesis that this population has been drawn from a Iognormal

population.

APPENDIX IV

APPENDIX IV

A COMPARISON OF COMPENSATED NEUTRON LOG POROSITIES OF THE
NONCONGLOMERITIC ZONES AND WHOLE-CORE POROSITIES OF THE
CONGLOMERITIC ZONES IN THE TRENTON FORMATION DOLOMITES

TABLE AIV- 1. COMPENSATED NEUTRON LOG POROSITV COMPARISON OF THE
NONCONGLOMERITIC ZONES IN THE TRENTON FORMATION DOLOMITES.

 

 

Sample Total Hicks-Martin H 1 Total Luck 2-12
(West) (East)
1 2.0 2.0
2 2.0 l .0
3 2.0 l .0
4 2.0 2.7
5 2.0 8.0
6 2.5 6.0
7 3.0 8.0
8 4.0 8.0
9 4.0 7.5
10 5.0 6.0
I 1 5.0 5.8
12 5.0 5.0
I 3 l .0 4.0
14 1.0 3.0
15 1.5 2.0
16 l .5 4.5
1 7 2.0 2.8
18 2.3 0.5
19 2.5 0.2
20 2.5 I .0
2 l 2.0 0.2
22 2.0 0.3
23 1.0 0.3
24 I .0 0.2

153

154
Table AW 1 -1 (cont'd.)

 

25 1 .0 0.0
26 I .0 0.5
27 1 .0 2.0
28 1.0 2.5
29 1.0 l .0
30 0.5
3 l 0.5
32 1.0
33 2.0
34 2.0
35 1 .0
36 1.0
37 1.0
38 l .0
39 l .0
40 0.5
41 0.8
42 0.8
43 0.8
44 0.8
45 0.8
46 0.8
Mean = 1.7 2.9
SD. : 1 .2 2.7

Appling a standard T-test to the data indicates that there is no significant
difference In porosity between the two populations.

155

TABLE AIV-2. WHOLE CORE POROSITY COMPARISON OF THE CONGLOMER ITI C
HORIZON IN THE TRENTON FORMATION DOLOMITES.

 

 

Sample Total Emil Faist 2-12 Total Luck 1-12
(West) (East)
I 1.7 3.5
2 1.1 5.3
3 1.7 4.5
4 2.6 6.4
5 2.7 3.7
6 3.7 4.3
7 1.9 2.7
8 2.1 2.9
9 2.3 1.8
l 0 2.8 1.6
1 I 2.3 8.9
12 1.6 5.9
13 2.7 6.9
14 2.0 3.8
15 2.7 3.1
16 1.4 3.0
17 3.8 6.2
18 2.1 8.3
19 2.8 6.4
20 4.2 2.9
21 I. l 3.2
22 3. l 3.4
23 3.5 3.7
24 6.1 3.6
25 0.9 3.5
26 2.2 3.9
27 3.0 2.7
28 1.8 3.7
29 2.6 4.0
30 2.0 3.3
3 l l. 1 4.8
32 3.4 3.3

33 2.5 3.6

Table AlV-2 (cont'd.)

34
35
36
37
38
39
4o
41
42
43
44
45
45
47
48
49
50
5 1
52
53
54
55
56
57
58
59
60
6 1
62
63
64
65
66
67
68
69
70
7 1
72

2.6
7.1
4.0
3.2
3.1
2.1
1.3
2.3
2.3
5.3
2.1
2. 1
6.3
4. 1
4.3
2.0
6.7
4.5
2. 1
2.5
4.7
3.8
2.1
3.0
2.3

156

3.2
3.3
2.4
3.6
3.4
3.0
4.0
2.5
2.7
2.9
4. 1
3.5
3.8
2.3
3.0
3.7
4.8
3.4
2.3
3.2
3.0
3.0
3. 1
3.0
2.7
3.4
3.4
5.4
3.4
3.0
4.5
5. 1
3.8
3.6
2.5
3.0
2. 1
2.3
2.6

157
Table AlV-2 (cont'd.)

 

73 2.5
74 3.1
Ilean : 2.9 3.7
5.0. = 1.4 1.4

Appling a standard T-test to the data indicates that there is no significant
difference in porosity between the two populations.

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