Clay mineralogy of the Willcox Playa and its drainage basin, Cochise County, Arizona
Item Type text; Dissertation-Reproduction (electronic)
Authors Pipkin, Bernard W.
Publisher The University of Arizona.
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Link to Item http://hdl.handle.net/10150/565620 CLAY MINERALOGY OF THE WILLCOX PLAYA
AND ITS DRAINAGE BASIN, COCHISE
COUNTY, ARIZONA
Bernard W/ Pipkin
A Dissertation Submitted to the Faculty of the
DEPARTMENT OF GEOLOGY
In Partial Fulfillment of the Requirements For the Degree of
DOCTOR OF PHILOSOPHY
In the Graduate College
THE UNIVERSITY OF ARIZONA
1964 THE UNIVERSITY OF ARIZONA
GRADUATE COLLEGE
I hereby recommend that this dissertation prepared under my direction by Bernard W. Pipkin entitled MClay Mineralogy of the Willcox Playa and its Drainage Basin, Cochise County, Arizona'* be accepted as fulfilling the dissertation requirement of the degree of Doctor of Philosophy.
______ss e/tat ion Director DateFa7^ /
After inspection of the dissertation, the following members of the Final Examination Committee concur in its approval and recommend its acceptance:*
♦This approval and acceptance is contingent on the candidate's adequate performance and defense of this dissertation at the final oral examina tion. The inclusion of this sheet bound into the library copy of the dissertation is evidence of satisfactory performance at the final ex amination. STATEMENT BY AUTHOR
This dissertation has been submitted in partial fulfillment of requirements for an advanced degree at The University of Arizona and is deposited in the University Library to be made available to borrowers under rules of the Library.
Brief quotations from this dissertation are allowable without special permission, provided that accurate acknowledgment of source is made. Requests for permission for extended quotation from or re production of this manuscript in whole or in part may be granted by the head of the major department or the Dean of the Graduate College when in his judgment the proposed use of the material is in the interests of scholarship. In all other instances, however, permission must be obtained from the author.
SIGNED: AC KNOWLEDGME NTS
The author wishes to express his sincere appreciation to Dr.
Joseph F. Schreiber, J r., who directed this study and provided the
core material and many soil samples for analysis. Professor Richard
Sloane generously took the electron micrographs and aided the writer
in their interpretation. Dr. Joseph F. Schreiber, Jr., Dr. John W.
Anthony, Dr. Spencer Titley, and Professor Richard Sloane critically
read the manuscript and greatly aided in its presentation. Dr. W. H.
Bradley of the University of Texas provided assistance in the quanti
tative aspects of clay-mineral determinations, and Dr. Thomas Bates
of the Pennsylvania State University offered helpful suggestions on in
terpreting electron micrographs. To the foregoing individuals, and to
the many graduate students and faculty members who aided the writer,
I express my sincere gratitude. However, the writer takes full re
sponsibility for all interpretations and conclusions contained within this
paper.
A portion of this work was done while the author was a National
Science Foundation Science Faculty Fellow at the University. Further
financial aid was provided through the College of Mines B. S. Butler
Scholarship and through National Science Foundation Research Grant
G-23746. iii Finally, I would like to thank my wife, Faye, for her patience, understanding, and encouragement during the course of this research. TABLE OF CONTENTS
Page
LIST OF ILLUSTRATIONS...... ix
LIST OF TABLES...... xi
ABSTRACT...... xii
INTRODUCTION...... 1
Location Climate and Vegetation ...... Purpose of the Investigation .. Scope of the Investigation...... Coring and Sampling Procedure
REGIONAL GEOLOGY......
Physiographic Setting ...... Stratigraphy......
General ...... Igneous rocks ......
Metamorphic rocks ...... CO CO OO 00 - J - ] tfr, GO H* M Sedimentary rocks ...... 11 Basin f i l l ...... 11
METHOD OF STUDY___ .’...... 13
Separation Technique ...... 13
Core samples ...... 13 Soil s a m p le s ...... 16
Preparation for X-Ray A n aly sis...... 18
Oriented aggregates ...... 18 Calcium-ion saturation ...... 20 Ethylene glycol adsorption...... 21 v vi
TABLE OF CONTENTS—Continued
Page
Therm al tre a tm e n t...... 22
Instrumentation...... 24
CLAY MINERALS ...... 28
General Statement ...... 28 Illite ...... 31 Montmorillonite 35 Vermiculite...... ; ...... 41 Kaolin M inerals ...... 45
Kao Un ite...... 45 H alloysite...... 49
Chlorite ...... 51 Interstratified M inerals ...... 53 Semi-Quantitative Analysis .'...... 58 Chemical Composition by X-Ray Analysis ...... 64
MINERALS OF THE WILLCOX PLAYA...... 69
Core M a te ria l...... 69
H lite...... 69 * Montmorillonite .* ...... 73 Vermiculite ...... 77 Chlorite ...... 78 Kao Unite...... 78 Mixed-layer m inerals ...... 79 Analcime..;...... 80 C alcite...... 85 Q u a rtz ...... 86 F e l d s p a r ...... 87 Heavy m inerals ...... ; ...... * 87 M in eralo id ...... 89
Electron M icro g rap h s...... 90
Method ...... 90 D isc u ssio n ...... 90 vii
TABLE OF CONTENTS—Continued
Page
Unknown m in e ra l...... 94
Playa S u rface...... 95 Soils and Alluvial Samples ...... 96
General ...... 96 M ineralo g y...... 96
ORIGIN OF PLAYA SEDIMENTS...... 102
Detrital Minerals ...... 102 Authigenic Minerals ...... 104
A nalcim e...... 104
Natural occurrences ...... 105 S ynthesis...... 110 Origin ...... I l l
.Calcite ...... 119 Evaporites ...1...... '...... 120
GEOLOGIC SIGNIFICANCE...... 122
Climatic Episodes...... 122 Diagenesis ...... 128
. Definition ...... 128 Diagenesis of clay minerals ...... 130 Willcox Playa sediments ...... 132 Paragenesis ...... 134
Mineral Facies ...... 135
SUMMARY OF CONCLUSIONS...... 140
a p p e n d ix a — contribution o f v a r io u s p a r t s OF MICA LAYER TO THE STRUCTURE FACTOR OF BASAL REFLECTIONS...... 142 TABLE OF CONTENTS—Continued
Page
APPENDIX B — HEAVY MINERALS OF THE C O R E...... 144
LITERATURE CITED ...... 146 LIST OF ILLUSTRATIONS
Figure , . » . . - . , ...... » Page
1. Location map of Willcox Playa ...... 2
2. Generalized geologic map of the Willcox Basin ...... 10
3. Flow diagram of analytical procedure ...... 27
4. Position and relative intensities of basal spacing of montmorillonite...... 40
.5. Relative intensities and basal spacings of oriented clay-mineral aggregates...... 43
6. Relative basal intensity curve of 2:1 dioctahedral clay m in e ra ls ...... 47
7. Effect of heating on diffractograms of kaolinite ...... 47
8. Curves of basal peak migrations for illite- montmorillonite mixed layers ...... 56
9. Mering's solution for random mixed-layer illite- montmor illo n ite ...... 57
10. Change in intensities for dioctahedral micas ...... 67
11. Clay-mineral composition of the core and playa surface ...... •...... 70
12. Diffractometer traces of core sample 46, depth 36 feet... 74
13. Diffractograms of analcime from Willcox Playa and pure analcime from Nova S cotia ...... 81
14. Analcime content of two core samples by the method of known additions...... 84 . X
Figure Page
15. Clay mineral distribution in soils, stream channels, and the playa surface ...... 97
16. Clay mineralogy, pollen zones, and Eh-pH conditions in the c o r e ...... 123
17. Diagenetic geochemical environment of the Willcox Playa se d im e n ts...... 138
Plate
1. Electron micrographs of the core material from a depth of three feet ...... 91
2. Electron micrographs of core material from a depth of 90 f e e t ...... 93 LIST OF TABLES
Table . Page
1. . X-ray diffractometer spacings and relative intensities for illite ...... 72
2. Common authigenic zeolite m inerals ...... 106
xi ABSTRACT
The Willcox Basin in southeastern Arizona is a closed drain age system some 1, 600 square miles in area. All drainage within the basin flows toward the Willcox Playa, which is the dessicated vestige of a Pleistocene lake named Lake Cochise by O. E. Meinzer. The mineralogy of a 140-foot core taken from the center of the playa was determined by X-ray diffraction, electron microscope, and polarizing microscope techniques. The clay mineralogy of the playa surface and of clay materials within the drainage basin was also determined and compared. The purpose of this work is to extend our knowledge of the genesis of clay minerals in an arid climate and of their diagenesis in a non-marine environment.
The core consisted almost wholly of a black mud with a siltzclay of 30:70. The bulk of the argillaceous material (-2 micron) was illite, with montmorillonite, mixed-layer illite-montmorillonite, and vermic- ulite in decreasing order of abundance. Two distinct monoionic divalent
montmorillonite complexes occur with basal spacings of 14.2-14. 7A and
15.2-15.4A, respectively. A moderately well-crystallized trioctahedral vermiculite occurs in about half the core samples. Kaolinite and chlo
rite occur in only trace amounts, the kaolinite occurrences being xiii restricted to the top few feet of core material. Authigenic analcime is a ubiquitous constituent of the -2 micron fraction in amounts of approxi mately 10 percent by weight. In addition to detrital quartz, feldspar, and heavy minerals, the coarse sediment fraction contains authigenic calcite euhedra and crystal aggregates and ostracod valves. Significant amounts of evaporites occur only as crusts at the playa surface and support the conclusion that Lake Cochise was a moderately saline
(brackish) lake. Comparison of the clay minerals of the core to those of the source area indicates the former are entirely of detrital origin.
The clay mineralogy of the playa surface is dominated by illite with less than 5 percent expansible phases present. Illite is also the predominant mineral in the source area; however, significant amounts of montmorillonite, chlorite, and kaolinite are being generated in the drainage basin. Discrepancies between the clay minerals of the lake bed and those of its source are best explained by selective removal of finer colloidal material at the playa surface by downward percolating w aters.
The origin of authigenic zeolites is reviewed, and data are presented that virtually exclude reorganization of volcanic glass or direct precipitation of an aluminosilicate gel to form analcime. It is believed that the diagenetic destruction of kaolinite to form analcime took place for the following reasons: (1) kaolinite is unstable in an alkaline environment, and all but a few of the 200 pH values obtained xiv fell between pH 9.0-9.5; (2) kaolinite is forming in minor amounts in the source area today; and (3) high pine pollen counts in the core interval
6-69 feet are indicative of higher precipitation and conditions more favorable for kaolinite genesis, yet this mineral is absent from the core below a depth of a few feet. A chemical mechanism is proposed to ac
complish the reaction kaolinite---- > analcime.
Synthesis of pollen data of other workers and the nonclay min
eralogy of the core indicates that the 140-foot sediment penetrated rep
resents one pluvial and one warm-dry episode of Pleistocene time.
Radiocarbon dates on carbonate from a depth of a few feet are greater
than 20, 000 years B. P ., and it is not known whether the inferred cli
matic episodes represent the first Wisconsin glacial advance and the
Illinoisan interglacial time, or whether they were affected during a
Wisconsin stade and interstade. It is concluded that the clay minerals
of the Willcox Playa core are of little use in paleoclimatic interpreta
tion. INTRODUCTION
Location
Willcox Playa is in southeastern Arizona just south of the town of Willcox (fig. 1). It occupies about 50 square miles at the southern
end of the Willcox Basin, a broad alluvial-filled valley without external drainage. The basin covers 1, 600 square miles from the drainage
divide at Aravaipa Creek on the north, to the divide at Pearce on the
south, and comprises the northern part of a larger valley known as
Sulphur Springs Valley. All the surface waters in the northern part of the valley drain toward the playa which is the dry lake bed of pluvial
Lake Cochise. At its maximum extent, Lake Cochise was 20 miles
long, 11 miles wide, and covered 120 square miles, or about 7 percent
of its drainage area (Meinzer and Kelton, 1913).
Climate and Vegetation
At Willcox the mean annual temperature is 89°F, the mean
annual rainfall is 18 inches, and the mean annual evaporation is 61
inches from a theoretical lake surface (Martin, 1963, p. 436). Summer
thundershowers release most of the precipitation, and the western part
1 2
ARIZONA
PHOENIX
YUMA
WILLCOX TUCSON
DOUGLAS
Figure 1. - -Location map of Willcox Playa. 3 of the playa is usually covered by a thin sheet of water after these tor rential rains.
The vegetation of Sulphur Springs Valley is characterized by desert grasslands and shrubs. The forms extant in the valley are sim ilar to those in southern New Mexico and northern Chihuahua (Martin,
1963) and are characteristic of the summer monsoon regimes of the
Southwest.
Purpose of the Investigation
A 140-foot core was collected in December, 1961, from near the center of the playa by the Geochronology Laboratory of the Univer sity of Arizona. The core was taken for palynological study to extend the known pollen stratigraphy of southern Arizona, and to yield informa tion regarding the effects of Pleistocene expansion of polar ice caps on the climate of the Southwest. The results of this work have been pub lished by Martin (1963).
The clay fraction of the 140-foot core contains a high propor tion of minerals formed by surface geological processes in a relatively
small, closed drainage basin. Much of this material was deposited in the playa area during a span of geologic time in which climatic condi
tions, and therefore the principal determinant on clay mineral genesis,
were different than they are today. The purpose of this work is to ex
tend our knowledge of the conditions under which fine-grained, 4 low-temperature mineral associations form and may be subsequently altered by different physiochemical conditions imposed upon them at the site of deposition. If late Pleistocene climates were drastically different than the present day, this difference should be reflected in the clay mineralogy of the playa. Thus clay minerals could provide a subsidiary tool for paleoclimatic determinations. The waters of Lake
Cochise were brackish, and if progress toward chemical equilibrium in such an environment is rapid, climatic changes would not be mani fest in the minerals of the core. In such a case, much can be learned about the diagenesis of clay minerals in a non-marine environment.
Scope of the Investigation
This study is a detailed analysis of the minus two (-2) micron fraction of the core, playa surface, soils, and alluvial sediments sur rounding the playa. The sand, silt, and heavy mineral fractions were
examined and described where necessary to interpret the minerals of
the clay fraction. The petrography of the coarser material will be de
scribed elsewhere.
Coring and Sampling Procedure
The core was taken by a commerical drilling firm utilizing a
5-foot long core barrel with a soil adapter. A 4-inch core was taken to
a depth of 74.1 feet and a 2.75-inch core from 74.1-140 feet. As the 5 core slug was removed from the barrel it was wrapped in mylar and sealed in plastic bags. The core was wire-split and described for lithology and color by Dr. Joseph F. Schreiber, Jr. Samples for size analysis and mineral content were taken from the inside of the core to avoid contamination from drilling fluids. The author obtained samples for analysis from air-dried material and from wrapped sections of the core that still retained interstitial water.
Electrode potentials and pH were measured with a Beckman
Zeromatic pH meter using a Beckman 39170 fiber-type calomel refer ence electrode, a Beckman 41263 glass pH electrode, and a platinum electrode for millivolt potential determinations (Mr. Austin Long,
University of Arizona). Mr Long obtained the values from a freshly opened core surface rather than a slurry. Reproducibility was to within
0.1 pH unit and 10 millivolts electrode potential.
Sediment samples from the playa surface were taken on a grid of 1 mile intervals and represent the top few inches of sediment. All well-defined stream channels around the playa were sampled with the thought that the clay in them contains representative minerals of the regolith in their respective drainage areas. Also they are representa tive of the sediment in the entire drainage basin being carried toward the playa.
Soils were sampled wherever good profiles were exposed free from agricultural influences. These were taken close to the surface 6 with emphasis on a high proportion of clay rather than selecting a con sistent zone in the profile. Although this method may not yield minerals characteristic of the soil, it is the surface material that is most subject to sheet erosion and thus carried into streams. REGIONAL GEOLOGY
Physiographic Setting
Sulphur Springs Valley lies in the mountain region of the Basin and Range province (Wilson, 1962). In this part of Arizona the prov ince is characterized by northeastward-trending mountain ranges sep arated by broad debris-filled valleys. The Willcox Basin is bounded on the east by the Pinaleno, Dos Cabezas, and Chiricahua Mountains and on the west by the Winchester, Little Dragoon, and Dragoon Mountains.
The valley floor is at an elevation of 4,135 feet at the playa and rises to 4, 500 feet at the northern divide. Mountains to the west reach eleva tions of 7, 500 feet, whereas the eastern ranges are higher with eleva tions of over 9,000 f^et in the Pinalenos. Rainfall is greater on the
eastern ranges (Coates, 1952), and this combined with the higher eleva
tions has produced much broader piedmonts fronting these mountains.
The net result has been a displacement of the axis of the Sulphur Springs
Valley toward the west. Sediment production from the Chiricahuas has
maintained the drainage divide near Pearce and in the past caused the
basin of Lake Cochise to migrate northward (Meinzer and Kelton, 1913).
Several geomorphic features related to the playa are worthy of
mention. Low beach ridges formed by the waves of Lake Cochise may 7 8 be found on the east and west sides of the playa. One well-defined ridge follows the 4 ,175-foot contour indicating a relatively long stand of the lake some 40 feet higher than the present dry lake bed. Cooper (1960) mapped a short ridge on the west side of the valley 60 feet above the playa surface. Thus, one can reckon pluvial Lake Cochise with depths of 40-50 feet if allowance is made for compaction of sediments, local subsidence, and deflation of the present lake surface.
Another striking feature is the large area of sand and clay dunes in the northeast quadrant of the playa, some of which reach heights of 50 feet. These are largely post-pluvial accumulations formed from coarse shoreline detritus of Lake Cochise. However, that the present flat is a deflation surface is shown in photographs by / Meinzer of low lake-clay buttes on the playa protected by vegetation.
The reduction of the playa surface is not known but is certainly on the order of a few feet.
Stratigraphy
General
It serves the purpose of this study to deviate from the usual chronologic description of rock units and group them according to ,1 chemical similarity and origin. Rocks exposed in the region range in age from Precambrian to Recent with all major chemical groups 9 represented except those found in mafic plutons. In compiling the ge ologic map (fig. 2), the writer has drawn liberally on the work of
Meinzer and Kelton (1913), Coates (1952), Gilluly (1956), Cooper (1960), and the County Geologic Maps of the Arizona Bureau of Mines.
Igneous rocks
Granitic igneous rocks include granite, quartz monzonite, and granodiorite high in alkalic feldspars and low in calcic feldspars and ferromagnesian minerals. Associated with these are hypabyssal rocks such as aplite, pegmatite, and rhyolite porphyry.
Large areas of effusive rocks mapped as rhyolite include rhy olite, rhyolitic tuffs, and agglomerates. They are the most common rock types and crop out over about 20 percent of the drainage basin.
Basalts, andesites, and related pyroclastics high in mafic constituents cover large areas in the Winchester, Galiuro, and Dos Cabezas Moun tains. Some of the basalts are very young and overlie Plio-Pleistocene alluvial deposits (Cooper, 1960; Pine, 1963).
Metamorphic rocks \
Except for local contact effects, metamorphic rocks in the region are all part of the Precambrian Pinal Schist. The formation is
composed of rhyolitic schist, greenstone, phyllite, metagraywacke,
and quartzite. Minor amounts of amphibolite are reported from the GENERALIZED GEOLOGIC MAP of the WILLCOX BASIN
EXPLANATION
Felsic Vdconics Granitic Intrusive Rocks
Mofic Vokxmics Metamorphic Rocks
Undifferentiated Granite Gneiss Volcanic*
Sedimentary Quaternary Rocks Lake Deposits
Principal Drainage
D rainage Divide
Figure 2. --Generalized geologic map of the Willcox Basin. 11
Dragoon Mountains. Granite gneiss underlies a large area in the north
eastern part of the drainage basin and supports the highest elevations
in the region.
Sedimentary rocks
Quartzite, sandstone, limestone, and shale occur in bands or
isolated pods throughout the drainage basin. Extensive erosion has re
moved much of the Mesozoic and Paleozoic strata, and they occupy but
5 percent of the drainage area (Coates, 1952).
Basin fill
Pliocene, Pleistocene, and Recent alluvial and lacustrine de
posits fill the valley and form bajadas fronting the mountain ranges.
Wells drilled in the valley have encountered alluvial gravels and sands
3, 000-5, 000 feet thick. Older alluvial deposits are well-cemented and
may form pediments on the flanks of the valley. Although the total
volume of basin fill is not known, it obviously represents a tremendous
reduction of the surrounding mountain ranges since the structural trough
was formed in mid-Tertiary time.
Lacustrine deposits occupy parts of the valley and probably
represent several episodes of sedimentation (Meinzer and Kelton, 1913).
At Willcox, the black clays of pluvial Lake Cochise were encountered at
an elevation of 3, 880 feet, whereas at Hado Station, at the north end of 12 the play a, lake muds were found at an elevation of 4, 070 feet. This in dicates that the lake occupied a much greater area and gradually con tracted to its present size as the climate changed. The total thickness of lake sediments is not known, but at least 200 feet of stiff black clay was reported in the Willcox well. The playa flat is 4,135 feet a. s. 1., and the black clay in the bottom of the Willcox well was at 3, 680 feet a. s. 1. This means that one would expect a minimum of 455 feet of lake deposits at the playa, as it occupies the lowest part of the Willcox
Basin. METHOD OF STUDY
Separation Technique
Core samples
Slices of core material were taken at about 5-foot intervals and air dried. A 10-15 gm sample was split onto filter paper, placed in a 100 ml, wide-mouth, screw-top jar and filled with tap water. The sample was then shaken vigorously by hand and allowed to slake for several hours or overnight. As a rule, the clays flocculated rather rapidly, and the salty supernatant liquid was siphoned off. More tap water was added, the sample resuspended, allowed to settle, and the clear supernatant liquid siphoned off. After the second or third wash ing in tap water, distilled water was added and the process repeated until a stable suspension was obtained. There are essentially two ob jectives in this procedure: (1) disaggregation, and (2) dispersion or deflocculation of the clay-size particles. Deflocculation is accomplished by removal of electrolyte in solution or by adding a polyanion dispers ing agent (Calgon). Disaggregation may be more difficult, and in those samples that have not completely disaggregated, one can see small clods or lumps of sediment. In the latter case the sample was placed in a 400
13 14 ml beaker with 200-300 ml of water and mechanically agitated in a rotary blender for several minutes. If small sample size is a problem, it is essential to control the amount of water used in washing and sus pending the sediment because dilute aqueous suspensions require exces sive evaporation (concentration) times.
The suspension was then poured into a 400 or 600 ml beaker, the water level raised to 6-7 cm above the bottom, and the sample agitated for several seconds in order to obtain a homogeneous disper sion at time zero. The beaker was allowed to stand for a time calcu lated from Stokes' Law for particles greater than 2 microns Stokes' sedimentation radius to settle. A rubber hose fitted with glass tubing on the ends was inserted to a depth of 5 cm and the clay suspension siphoned into a 250 ml beaker. The hose was washed with distilled water, inserted to a depth in the next beaker, the clay suspension siphoned off, and so on. It should be mentioned that the exact time of withdrawal is not critical as long as the elapsed time is sufficient to settle out the +2 micron particles. That is to say, the settling velocities of clay-size particles are so slow that siphoning off the clay suspension
even an hour after the prescribed settling time dictated by Stokes' Law
results in loss of only a small portion of the sample. Mackenzie (1956)
describes a method for handling large numbers of soil clay samples
which are sedimented overnight (5 p. m. to 9 a. m .) and withdrawn from
a depth of 10 cm. In this manner, depending on the temperature, the 15 particle-size will be some figure less than 2 microns. Some important aspects of the subject are further discussed in papers by Bray, Grim, and Kerr (1935); Krumbein and Pettijohn (1938); Grim and Johns (1954);
Hathaway (1956); and Brindley (1961a).
The material remaining in the bottom of the sedimentation vessel was screened through a 250-mesh sieve and the retained frac tion washed onto filter paper, air-dried, and stored for microscopic study. The material passing the screen was suspended, allowed to settle for a few minutes, and the clay and fine silt in suspension de canted. The decantation was repeated until a reasonably clean medium to coarse silt was obtained. This fraction was also washed onto filter paper, dried, and stored in a sample envelope.
The -2 micron separation was evaporated on a hot plate at
40°C to about one-half the original volume. As evaporation proceeds,
electrolyte concentration increases and flocculation may occur. In
such cases, the supernatant liquid was poured off and the clay resus pended in distilled water. When the clay slurry was sufficiently con
centrated to form a translucent clay cake on a glass slide, a portion was
stored in a 5-dram vial and the remainder was evaporated to dryness,
lightly crushed, and similarly stored.
Other methods of separation were tried, including suspension
with a polyanion dispersing agent (Calgon) and centrifugation. For
small samples of high clay content, the method described was found to 16 accomplish the desired separation without subjecting the clays to chem ical treatment before irradiation. Using the foregoing procedure one can process 10-15 samples simultaneously utilizing a minimum of laboratory space and actual time spent manipulating equipment.
Soil samples
Soil and stream-channel samples required a different separa tion technique because of their high percentage of organic matter and variable proportions of clay. Depending upon the clay content of the sample, a 25-50 gm split was put into a one quart, wide-mouth, screw- top jar and filled with tap water. The sample was shaken, allowed to
slake for several hours, and the supernatant liquid poured off, carrying with it discrete particles of organic matter that had floated to the sur face. The process was repeated several times or until most of the
visible organic matter had been decanted. Then the jar was half-filled
with distilled water to which was added 5-10 ml of 20 percent HgOg.
The sample was agitated with a rotary blender for about a minute and
allowed to stand from 30 to 60 minutes. As oxidation proceeded, an
occasional stirring with a glass rod brought more organic particles
into contact with the oxidizer. The carbon dioxide evolved bubbled to
the surface and brought with it much of the finer organic matter in the
soil. After the suspension had cleared and the production of carbon
dioxide had ceased, the supernatant liquid was siphoned off and the fine 17 particles of organic matter adhering to the jar at the meniscus were wiped off with a laboratory tissue. The writer found that complete oxidation of organic carbon in soils and alluvial samples required in ordinate amounts of HgOg and time. Therefore, at this stage of the separation, the sample was screened through a 250-mesh sieve and most of the remaining organic material and the coarser constituents of the soil were removed. Prior to screening, distilled water was added and the sample agitated for 10 minutes with a rotary blender. If lumps or balls of clay were noted during the sieving, the sample was stirred again in the blender and more H2 O2 was added. The presumption that films or coatings of organic matter cementing clay grains removed by this process was verified by soil-clay diffraction patterns exhibiting a high degree of preferred orientation indicated by strong basal reflec tions. Weak 00JZ reflections, particularly the reduction of basal peak intensities relative to prism or pinacoid reflections, indicates poor orientation of clay flakes on their basal planes. Such samples should be further treated with H2 O2 lest one draw erroneous conclusions re
garding degree of crystallinity or disordering of the clay.
The disaggregated, organic-matter free suspension was re
turned to the quart jar and 100 to 200 mg of dispersing agent was added
to it. The sediment was stirred and allowed to settle for a time calcu
lated from Stokes’ Law to obtain the -2 micron fraction. The top 5 cm
of suspension was siphoned into a 400 cc beaker; the slurry was usually 18 thick enough to prepare slides of oriented aggregates without additional
concentration. Some coarse-grained soils and some channel sands re quired concentration by evaporation at 40°C on a hot plate or in an oven.
The slurry and dried clay were stored in the same manner as the core
sam ples.
Preparation for X-Ray Analysis
Oriented aggregates
Slides of oriented aggregates (Bradley, Grim, and Clark, 1937)
were prepared by depositing the clay slurry from a glass dropping pipette onto a 27 by 46 millimeter petrographic slide and allowing it to
dry at room temperature. The result of such a preparation is that dif
fraction of X-rays is optimized along (00.fi) reciprocal lattice nodes
(Bradley, 1954) and minimized in other crystallographic directions.
Enough suspension was applied to the part of the slide to be irradiated
so that it was held at the edges of the glass by surface tension of the
liquid. In routine mineralogical analysis no attempt was made to
standardize the amount of clay X-rayed. Usually 6 to 8 drops of slurry
contained enough clay to give good basal reflections. Brindley (1961a)
states that slides prepared in this manner should contain about 15 mg
of dry mineral matter. It was found that a dry clay cake which is trans
lucent to transmitted light contained enough diffracting layers to produce 19 good basal intensities.
Several precautions should be taken to insure that the flat pow der layer produced has a uniform thickness and maximum preferential orientation. The glass slide should be clean and dry so that surface tension of the aqueous suspension will hold the slurry on the mount along the unwetted edges. A smooth, level surface is essential for clay grains to spread uniformly on the slide. A piece of plate glass, about 1-foot square, is an ideal substrate for mounting glass slides.
Finally, the slurry should be added a drop at a time from the center of the slide toward the edges, without stirring or disrupting the surface in any way. Intensities of basal reflections of samples that had been
stirred after mounting were inferior to those prepared by adding drops from the center outward. Thin powder layers contain too few planes
satisfying the Bragg equation to give unequivocal diffraction effects.
The addition of more slurry to the dry clay film only serves to disrupt
the orientation of the first layer. In such cases, the slide should be
cleaned with a razor blade and a new mount prepared with more clay or
a more concentrated suspension. Brindley (1961a) presents a good dis
cussion and summary of the literature on preparation of clay aggregates
for X-ray analysis.
In order to identify and characterize clay minerals, several
diffraction patterns of a specimen should be obtained. In actual prac
tice the procedure depends on the number of samples and how much 20 information the investigator wishes to obtain. With cation saturation, polar-liquid adsorption, acid dissolution, and thermal techniques, as many as a dozen patterns may be required (Brindley, 1961a). In this study five diffractograms of each sample were considered a minimum number to identify the mineral phases present. They were: (1) un- treated oriented aggregate, (2) Ca— saturated oriented aggregate, (3)
Ca™ slide treated with ethylene glycol, (4) oriented aggregate or glycolated slide heated to 190°C, and (5) oriented aggregate heated to
590°C. In addition to these, K—, Mg— Li—, and glycerol-saturated slides were prepared for more specialized purposes. Non-oriented powder specimens were prepared for the core samples in order to de termine the d(060) spacing.
Calcium-ion saturation
Two methods of calcium-ion saturation were used. Several samples were passed through an ion exchange resin (Amberlite-IR-120) that had been charged with a IN CaClg solution and rinsed with distilled water. This technique is used in the sedimentary petrography labora tory of the U.S. Geological Survey (Hathaway, 1956). However, the procedure is not practicable with small samples because of dilution resulting from washing the sample through the resin column. A second technique, and the one ultimately used on all samples, was to smear a drop or two of IN CaClg on the slide and add the clay slurry to this. 21
The X-ray results were as good as those obtained from clays passed through the exchange resin and preparation took much less time. There was the problem of a synthetic salt phase forming on the slide when
CaClg in excess of the amount needed to fill cation exchange sites in the clay occurred. This phase may be taken into account by comparing diffractograms of untreated and CaClg treated slides.
Ethylene glycol adsorption
The expansion of basal spacings of montmorillonite - type min erals with the introduction of water or other polar liquid between layers is a diagnostic feature of many clay species. Ethylene glycol and glycerol (glycerin) are most commonly used to convert expandable 2:1 layer silicates to the organic liquid "solvated" phase accompanied by expansion in the [001] direction. Clays imbibe glycol or glycerol readily so that the choice of solvation technique depends only upon how it affects orientation. Common methods are (1) mixing the clay min erals and liquid on the slide and air drying until the clay cake appears slightly moist, (2) applying the polar liquid by means of a dropper or capillary to the edges of a dry clay cake until adsorption is complete, and (3) glycol vapor adsorption at reduced pressures or at slightly ele vated temperatures. Methods (1) and (2) are simple and fast but seldom yield solvated slides without some disruption of orientation. Brunton's
(1955) vapor-pressure technique is based on the relatively low boiling 22 point of ethylene glycol and the adsorption of its vapor by clay films without wrinkling. A slightly modified version of Brunton's technique was used in this study. A 50 cc beaker with glycol-saturated cotton at the bottom is placed on a hot plate at about 50°C. The slide is placed sample-side down across the mouth of the beaker and a 250 cc beaker inverted over the smaller beaker in order to contain the vapor. The temperature should be adjusted so that enough vapor is produced to give a well-glycolated slide in one or two hours. Care should be ex ercised in working up to the optimum temperature because greatly in creased vapor pressure will cause droplets of glycol to condense on the
slide and ruin orientation. Vapor adsorption may also be effected at
low temperature in a vacuum. However, complete solvation takes 12-
24 hours and only a few slides may be treated at one time. When the
slide appeared moist with glycol, it was blotted with a laboratory tissue
and stored in an aluminum dessicator containing some of the liquid.
Kunze (1955) points out that clays can lose ethylene glycol at room tem
perature, and a contraction of basal spacings ensues which may be mis
interpreted as mixed layers. For this reason, all specimens were
X-rayed within a few hours after removal from glycol vapor.
Thermal treatment
Diffractometer patterns of heat-treated specimens, when com
pared to patterns of their untreated equivalents, can yield valuable 23 information concerning the existence of thermally unstable phases in clay-mineral mixtures. One slide was placed in a pre-heated oven at
1909C for at least 2 hours and X-rayed before cooling in a semi-con trolled atmosphere. This procedure collapses the expandable com ponents and gives a measure of their relative abundance. Humidity control is necessary because rehydration of montmorillonoids is rapid in moist air (Milne and Warshaw, 1956; Walker, 1961). Air was pumped through a column of dessicant (Drierite) and directed onto the slide surface by a rubber hose introduced through the slotted scatter shield of the diffractometer. Although this method is crude, a high flow of air was maintained which kept the air surrounding the slide at less than 20 percent relative humidity, whereas the laboratory air was consistently two or three times that value. A more sophisticated and reliable adaption of diffractometer equipment for control of clay speci mens is described by Milne and Warshaw (1956)e
A second slide was heated at 590° + 10°C for 2 hours and air- quenched. The changes that occur to clay minerals at this temperature are discussed in another section of this report. Glass slides at this temperature will warp or break when cooled rapidly. If the slide is carefully withdrawn from the oven and allowed to cool on a pre-heated flat surface for a few minutes, warping is negligible. Warshaw and Roy
(1961) suggest the use of Vycor slides for high-tem perature investiga tions. 24
Instrumentation
All patterns were recorded by means of a Norelco X-ray Re cording Diffractometer using nickel-filtered copper K& radiation ( X =
1.5418A). The X-ray tube power was 40 kv, 20 ma; the detector oper ating voltage plateau was determined in the manner suggested by Klug and Alexander (1954, p. 262); and the amplifier gain was 45-50. Di vergence and scatter slits were kept at 1° 20, and the receiving slit width at 0.003 inches. Wider slits were tried on oriented specimens in the high-angle range (450-65° 20, but the loss of resolution was not compensated by the increased intensity. The untreated oriented aggre-
i Q gate, glycol-solvated, and Ca treated specimens were scanned at
1/4° 20 per minute, from 2o-30° 20, with typical settings of time con stant = 16, counting rate = 100; time constant = 8, and counting rate =
200. At faster scanning speeds the slow time constant presents a prob lem of resolution between closely spaced peaks in the low 0 20 region;
however, at faster pulse response times the background noise is so ob
jectionable that some resolution loss can be tolerated. When the product
of the (time constant) x (scanning speed) is four or more, there is con
siderable reduction of peak intensity, line shift in the scanning direction,
and asymmetrical broadening of peaks (Parrish, 1960). Fortunately,
the percentage reduction of peak intensity is the same for all lines
scanned in the same direction; therefore, relative peak intensities may 25 be safely used (Parrish, 1960). Also, the use of heat and chemically treated slides usually clarified any doubt regarding real or imaginary diffraction peaks at low angles. Heat treated core specimens were run at 1° 20/minute through the angular range 20-15° 20. The faster speed was used partly to prevent rehydration and partly because the informa tion yielded by thermal treatment did not necessitate higher resolution.
Angular values of diffraction peaks were read directly from the pattern and then converted to d spacings in Angstrom units using the tabulation of Switzer et al. (1948).
Soil clays were irradiated with the same instrument settings as the core material, except when well-crystallized or nearly mono miner alic clay suites permitted faster scanning rates. All clay films were centered visually with respect to the axis of rotation of the gonio meter to insure the maximum number of reflecting planes in the path of the X-ray beam.
The convention is followed that Miller indices of crystal planes and faces are enclosed in parenthesis (hkl), and the notation for reflec tions from these planes are without parenthesis hkl. Older data re ported in Kx units are so indicated, but they were not converted to
Angstroms because the difference has no significance in this study.
Powder photographs were taken to give additional information about clays and accessory minerals and to corroborate diffractometer data. Both the 114 and 57 millimeter cameras were used with copper 26
K& radiation. As the 17A spacing of montmorillonite gives reflections within 5 mm of the direct beam in the 114 mm camera, photographic recording of clay minerals with an unmodified camera is not feasible.
However, powder photographs are valuable for recording weak reflec tions from planes with spacings less than 10A. The data from photo
graphic recordings are summarized in the body of the paper.
The analytical procedure is shown schematically in figure 3. SAMPLE PREPARATION PROCEDURE
15 gm split
Separation of Size Fractions
Heavy mineral < 2 / i Slurry S a n d a n aly sis
Random Powder Pow der O ptical 060 reflection Photographs S tu d y
Oriented Aggregate B inocular slides
E thylene I 9 0 " C 5 9 0 * C Glycol 2 hours 2 h o u rs Figure 3. --Flow diagram of analytical procedure. t-3 o CLAY MINERALS
General Statement
In modern technology the term "clay" may denote the firing characteristics, plasticity, or grain-size of a rock or sediment. To the ceramicist clay implies a material that becomes plastic when wet and rock-hard upon firing, whereas the geologist, engineer, and soil scientist are most likely to concede with Grim (1953) that: "In gen eral, the term clay implies a natural, earthy, fine-grained material which develops plasticity when mixed with a limited amount of w ater."
Mackenzie (1963) notes that Grim's definition is that proposed 400 years earlier by Agricola rendered in modern idiom. One of the difficulties in definition is that some clays are not plastic, the so-called "flint- clays" for example (Grim, 1953), and many rock-forming minerals develop plasticity when reduced to extremely fine particle size, e. g. quartz (Mackenzie, 1963). A more serious obstacle to the acceptance of a single definition lies in the use of clay as a particle size term.
United States geologists designate as clay that material which is finer than 4 m icrons in diam eter in accordance with Wentworth*s (1922) size grades. Engineers understand clay as the minus 200 sieve material
28 29
(74 microns) which has a high plastic index relative to the liquid limit, i. e., high plasticity (Corps of Engineers, 1953), and the U.S. Bureau of Soils considers clay as a soil consisting of grains with diameters smaller than 50 microns (Howell, 1960). A definition based on grain- size is unsatisfactory because of the different standards used and be cause many commercial clay deposits are not clays by these definitions
(Pettijohn, 1957). However, if a mineralogical restriction is placed on clay, then there is no general term for very fine-grained clay-grade sediments in the Wentworth grade scale. Clay in this work is in the sense of Grim's (1953) definition, with the added qualification of a maximum diameter of 4 microns.
Just as the term "clay" is indefinite, so is the meaning of clay mineral. Any mineral found in clay is not a clay mineral because the definition would embrace a multitude of accessory constituents which contribute little to the physiochemical characteristics of clays. Also, the vagueness of the term "clay" militates against such a definition
(Mackenzie, 1963). Clay minerals are frequently referred to as hydrous aluminum silicates. Such references preclude minerals low in alumina, such as nontronite, saponite, and hectorite, which are normally con sidered in clay-mineral studies (Ross and Hendricks, 1945). Grim
(1953) anticipated the low alumina minerals when he defined clay min erals as ". . . . essentially hydrous aluminum silicates with magnesi um or iron proxying wholly or in part for aluminum in some minerals 30 and with alkalies and alkaline earths present as essential constituents in some of them. ” The lack of structural criteria in the definition is viewed as an important omission by the writer because hydrous alumino silicates could just as well be zeolites as clay minerals. Mackenzie
(1959) places emphasis on crystal structure and states: "Crystalline clay minerals are hydrated silicates with layer or chain lattices con sisting of sheets of silica tetrahedra arranged in hexagonal form, con densed with octahedral sheets; they are usually of small particle size. ”
In a later work on classification and nomenclature he states that objec tions were raised because of the exclusion from the definition of amor phous clay material and the physical property of plasticity (Mackenzie,
1963). He states that in view of the fact that clay minerals belong to one or another of the ordinary mineral species, it is probably better to give a very general definition such as "the minerals which normally predominate in the colloidal fraction of clay, rocks, soils, etc." The word "normally" was inserted to exclude nonclay mineral impurities occurring in all clays.
In this work the term clay mineral is intended for fine-grained hydrous silicates with layer or chain structures, most commonly con sisting of two tetrahedral layers enclosing one octahedral layer (2:1 mica type); and one tetrahedral layer and one octahedral layer (1:1 kaolin type); and including stacking polytypes and randomly stacked
(disordered) structures. Clay mineral species may be produced by 31 the regular interstratification of different kinds of layers (chlorite, corrensite, rectorite, etc.) and mixed-layer assemblages by random interstratification of layers of different types.
Nonclay minerals are those which fall outside the foregoing definitions for reasons of chemistry, structure, or grain-size. These most commonly include calcite, quartz, dolomite, feldspar, pyrite, mica minerals, sulfates, cristobalite, tridymite, and gibbsite. There is probably no clay which does not contain some nonclay mineral coarser than the clay fraction (Grim, 1953). Fortunately, most of the nonclay minerals are relatively coarse and easily separated from the -2 micron fraction. It is now common practice in laboratories concerned with clay mineral determinations to separate the clay fraction at 2 microns or less. The loss in clay minerals is more than compensated by the re duced interference from nonclay minerals. Grim (1953) argues that for this reason there is a fundamental reason for placing the upper limit of the clay fraction at 2 microns.
Elite
The name "Elite" was proposed by Grim, Bray, and Bradley
(1937) as a non-specific term for the clay minerals of rocks and sedi ments belonging to the mica group. The type iUite from Elinois shales has less KgO, m ore H2 O, more mafic components, and a slightly 32 smaller cell (1/3 percent) than muscovite. It may be represented by a structural form ula of the type
Kl - x+y(A1> Fe+3)2„y (Mg, Fe+2)ySi3 +xAI1„xO10(OH)2 where x ranges from . 3-. 6, and x greater than y. A similar mineral had been reported by many investigators prior to 1937 and considerable
synonomy exists, e. g., hydromica, hydrous mica, clay mica, and
glimmerton to name a few. Ross and Hendricks (1945) maintained that
the name "bravaisite” has priority over illite for a "mica-like mineral with about half the potassium of muscovite." Grim and Rowland (1942)
concluded that type bravaisite is a mixture of illite with a little mont-
morillonite and therefore not a valid species.
Clay mineral investigations since the 1940's have done much to
reconcile chemical analyses and diffraction characteristics of illites as
they relate to micas. Structural similarities between montmorillonite
and mica permit inter layering of the two types of layers within a single
coherently scattering crystallite. Loss of potassium, increased hydra
tion, and asymmetry of the 10A reflection are consequences of such
mixed layering. Thus, Bradley (1945) demonstrated that bravaisite is
not a mixture of minerals but rather a mixed-layer assemblage within
the illite-montmorillonite range. It now seems quite likely that a com
plete series exists between dominantly hydrated expandable layers and
non-expanding illite (Grim, 1953; Weaver, 1956; Jonas, 1960; Bradley 33 and Grim, 1961). For this reason Weaver (1959) defined illite as a 10A clay mineral with less than 10 percent expandable layers.
Experimental studies on micas at the Geophysical Laboratory of the Carnegie Institution (Yoder and Eugster, 1955; Smith and Yoder,
1956) reviewed by Yoder (1959) have focused attention on polymorphism of mica crystallization as related to the physical and chemical condi tions of their formation. Polymorphism, which involves identical two- dimensional layers stacked in different ways, is commonly described as polytypism (Azaroff, 1960, p. 364). Silicon carbide and ZnS are classic examples, and the clay minerals are quite analogous to these.
Since the possibly ways of stacking hexagonal closest-packed layers are legion, it is impractical to name polymorphic varieties as is done with the SiOg minerals. Rather, a system of notation is used where a nu
merical coefficient indicates the number of basic structural units in the
cell and a letter denotes the symmetry of the cell. Thus muscovite-
1M is one structural layer (10A (001) periodicity) with monoclinic sym
metry, and muscovite-2M consists of two structural layers (20A (001)
periodicity) with monoclinic symmetry. Yoder and Eugster (1955) have
shown that for synthetic muscovite and paragonite, the structural trans
formation proceeds IMd (random stacking) to 1M to 2M%. They hold
that the 2M% polytype in sediments indicates a detrital origin because
of the higher temperature of synthesis. They found that the IMd and
1M muscovite were common in materials called illite, sericite, 34 hydromica, secondary mica, and glimmer ton. Yoder (1959) states that illite M. . . . consists of mixed layers of IMd and 1M muscovite and montmorillonite and may contain 2Mj or 3T muscovite as an additional phase.” He cautions against bulk composition as evidence for solid solution series muscovite-montmorillonite on the grounds that the limits of substitution of Si+4 for KtAl+3 have not been established. Yoder and his co-workers favor the position that illite should be used only as a field term, and that after further laboratory analysis the mica poly morph and degree of mixed layering should be specified.
Several circumstances militate against such an ideal analytical procedure. It is often impossible to determine the mica polytype in very fine-grained clays and clay mixtures. Secondly, X-rays cannot differentiate IMd polytypes formed by degradation of better crystallized material from those formed by improved crystallization of less mica ceous matter (K-bentonites). Genetically, there are two IMd polytypes, one of which is unlike the synthetic product (Bradley and Grim, 1961).
Finally, illite is a good descriptive term for those montmorillonite-
illite mixed-layer assemblages that do not depart significantly from a
10A basal spacing (Bradley and Grim, 1961).
Illite is used here for a dioctahedral clay mineral exhibiting
an integral sequence of basal reflections beginning at 10A„ No change
in basal spacings is affected by treatment with ethylene glycol or heat
up to about 450°C. Inter layering of expandable units is indicated by a /
35 pronounced tailing off of the 8.8° 20 peak toward the low-angle side and of the 26.6° 20 peak toward the higher angles. Heat treatment and ir radiation under controlled humidity conditions yields a sharper and more intense 10A reflection due to scattering contributions from col lapsed hydrated phases. Hlite is not synonymous with a particular clay- mica polytype and therefore the 10A spacing is referred to (001), even though this index is forbidden by space group in the very common mus- covite-2Mj.
Montmorillonite
In contemporary usage, montmorillonite refers to both a group of clay minerals that expand parallel to [001 ] (except vermiculite) and to a specific clay mineral (Ross and Hendricks, 1945). Like the micas, the basic structure is that of pyrophyUite-talc with appropriate ionic substitutions producing the physical and chemical properties unique to montmorillonites. Whereas micas have a high lattice charge due to proxying of Al+^ for Si+^ in tetrahedral coordination, montmorillonites have about one-third the charge imbalance seated in octahedral and/or tetrahedral sites. This low charge is compensated by monovalent or divalent (rarely trivalent) cations loosely held between the layers. The resultant mineral consists of randomly stacked structural units which are easily expanded by penetration of water or other polar liquid 36 between layers and which exhibit surprisingly constant cation exchange properties (Ross and Hendricks, 1945).
The group is conveniently divided into subgroups based on the cation content of the octahedral layer. Dioctahedral montmorillonites are represented by a solid solution series, montmorillonite-beidellite- nontronite, in which 2.0-2.2 of the three possible octahedral sites are occupied. Aluminum4"^ proxy Si"1"^ operates in beidellite and nontronite to bond the exchangeable external ions. In the nontronite -beidellite
series Fe+3 proxy Al+^ in 6-fold coordination is complete, whereas in
montmorillonite-beidellite Mg+2 proxy Al+^ is partial and limited to the
observed range of exchangeable cations. The series saponite-hectorite
is typical of montmorillonites in which the number of ions in octahedral
coordination is three. Saponite may be viewed as the magnesian analog
of beidellite, whereas hectorite is a special case in which limited Li+
, O proxy Mg in octahedral sites accounts for the lattice charge and ex
changeable cations.
Ross and Hendricks (1945) found no members intermediate in
octahedral site occupancy between dioctahedral and trioctahedral mont
morillonites. This agrees with the findings of Yoder and his coUeagues
that there is no known solid solution series between dioctahedral and
trioctahedral micas (Yoder, 1959). Structural formulas of the principal
members of the montmorillonite group are (after Ross and Hendricks,
1945, half-cell formulas): 37
(a) montmorillonite (All. 67 Mg. 33)Si4°10 (OH)2
(b) beidellite (AI2.17)813.17AI. 83 °10 (°h)2
(c) nontronite (Fe2)Si3. 67A1.33 Oiq (OH)2
(d) saponite (Mg3.00)Si3. 67A1. 33 °10 (°h)2 (e) hectorite (Mgg. 67 Li. 33)814 O10 (OH)2
Many other montmorillonite minerals have been described, and these are summarized by Ross (1960) and Mac Ewan (1961). Ross and
Hendricks (1945) point out: "Other elements such as Fe+%, Mn+2, Ni+^, would play a role identical with Mg+^, and separate formulas need not be written to represent them as they normally occur in very minor proportions."
The term "aluminian montmorillonite" to replace beidellite has been proposed (Ross, 1960), but the work of Weir and Greene-Kelly
(1962) points clearly to the existence of beidellite with a SitAl of about
3:2.
Several alternatives have been proposed to overcome the dual meaning of montmorillonite. The term "smectites" has priority and was widely used in mineralogical works of the 18th century (Ross and
Hendricks, 1945). It was proposed by the Clay Minerals Group (British) sub-committee on nomenclature as a substitute for the broad meaning of montmorillonite (Brown, 1955a). MacEwan (1951) suggested the term
"montmorillonoid" for minerals of the group. Both terms have met 38 resistance in the United States, and MacEwan himself does little more than mention the term in his summary of the montmorillonite minerals
(MacEwan, 1961). In spite of the practicality and justification of re vised nomenclature, montmorillonite is entrenched in the literature and its replacement is not imminent.
The oriented aggregates procedure in clay mineral studies
limits reflections almost exclusively to the class 00J2. . In the mont morillonite group, basal spacings vary according to (1) the interlayer
(exchangeable) cation, (2) the kind and amount of liquid adsorbed, and
(3) the nature of the particular mineral involved. Bradley, Grim, and
Clark (1937) showed a discrete series of hydrates for H-montmorillonite
with basal spacings (001) of 12.4A, 15.4A, 18.4A, and 21.4A. This
corresponds to the formation of water layers 1, 2, 3, and 4 molecules
thick, respectively. The stepwise hydration of montmorillonite is now
generally accepted, and the hydrate stable at a given vapor pressure
depends on the exchangeable cation. Thus Na-montmorillonite forms a
one-layer hydrate with a 12.4A (001) spacing, and Ca-montmorillonite
forms a double water layer with a basal spacing of about 15.4A at
humidities normally encountered in the laboratory. It is of utmost sig
nificance that a continuous change in d(001) with uptake of water is pos
sible. This results from random mixing of layers in different stages of
hydration or, as McAtee (1956) has shown, the random interstratifica
tion of Na-Ca-montmorillonites with one and two water layers, 39 respectively. MacEwan (1961) points out that: "Many measured spac- ings, especially in earlier work, are those of interstratified mixtures, and thus have no absolute significance."
Glycerol and ethylene glycol form two-layer complexes with
Na or C a-saturated montmorillonite and yield (001) spacings of about
17. 7A and 17.0A, respectively. Kunze (1955) has shown that m ontm oril- lonites may lose ethylene glycol at room temperature in as little as an hours time. He cautions that glycol-solvated mounts which have been exposed to the atmosphere could yield results erroneously interpreted as mixed layering.
All montmorillonites lose interlayer water at 100-200°C and collapse to d(001) 9.5-10.0A. The temperature at which this water loss is irreversible depends on the interlayer cations, and for Na or Ca- montmorillonites it lies between 300° and 400°C (Grim, 1953). This behavior provides a powerful identification tool because these tempera tures lie well below the OH structural-water endotherm for kaolinite,
illite, and chlorite. When clay mineral mixtures containing montmoril
lonite are dehydrated and X-rayed at low relative humidity (20 percent),
the addition of coherent scattering from the collapsed structure to the
10A peak of illite is a measure of the expandable component present.
It is apparent that the diagnostic property of montmorillonites
is induced shifts of basal spacings. The characteristic d-spacings are
illustrated in figure 4 and are the criteria by which this mineral was Figure 4 .-- .-- 4 Figure cr e o intensity. 10 9 8 7 6 5 6 7 8 9 0 1 M i ______If) O o 0) o —] 5 0 24 54 70 7 20 2 177 17.0 15.4 12.4 10 .5 9 ______------oiin n rltv itniis f aa saig f montmorillonite. of spacing basal of intensities relative and Position § 0 O T I I I I I I T 1 Mixed - Layer Layer Mixed - Region dA i______U ) O O i n t ^ l - i ♦ Z o ______M ixed-Layer ixed-Layer M Region l ______M M ♦ 4- O 5 o* ______i 1 0 ______O o £ >. i o 41 recognized in mixtures. Basal reflections are not diagnostic of specific members of the group and are useless in this respect. General reflec tions are more useful, and in the absence of interference the 060 reflec tion can distinguish the montmorillonite-beidellite-nontronite series from saponite and hectorite. Greene-Kelly (1953) has devised a test for
montmorillonite and beidellite, but one must first know that the assemblage belongs to this series. The Li-saturated specimen is heated overnight at 200o-300°C and then saturated with glycerol. Failure to expand in
dicates montmorillonite^ as both beidellite and nontr onite expand to 17, 7A
with interlayer glycerol,
Vermiculites have been reported which expand to about 17A
with glycol or glycerol. Care must be exercised in using (001) dilation
as the criterion for positive identification of the montmorillonite group.
Vermiculite
Analysis of macroscopic vermiculite from West Chester,
Pennsylvania, yielded the half-cell formula (after Walker, 1956):
(Mgl. 92 F e ^ g A1 2 2 Ti# n F e ^ g) (Si2. 72 A11.28> O10 (OH)2 . 38M g4.43 h 2°»
with a cation exchange capacity of 158 meq/100 gm. Recent work in
dicates that the difference in properties between clay vermiculites and
montmorillonites is essentially a difference in layer charge (Walker, 42
(1961). Vermiculite has almost twice the layer charge and thus double the cation exchange capacity as the mineral montmorillonite. High charge montmorillonites (Groschlattengrun saponite of Hofmann et al.,
1956) and low charge verm iculites have very sim ilar properties.
Warshaw and Roy (1961) hold that saponite-vermiculite probably form a
continuous series and should not be separated. They classify vermiculite
as a separate sub-group of expandable 2:1 layer silicates on the basis of
its failure to expand to 17A with glycol solvation. Similarly, Hofmann
et al. (1956) favor the concept of a continuous gradation between ex
pandable and non-expandable clay minerals, and they place a layer charge
value of 0. 55 equivalents per O^q (OH)2 unit as the boundary between
groups.
Vermiculite forms a stable two-layer hydrated phase with a
basal spacing of 14.36A. In theory it can be distinguished from chlorite
by a very strong 001 reflection (fig. 5) and collapse to 10A with moder
ate heat'; and from montmorillonite by its inability to form double-layer
complexes with glycol or glycerol and expand to 17A. Walker (1957)
reduced four macroscopic Mg-vermiculites to clay-size and found no
appreciable displacement of d(001) with ethylene glycol. Replacement
of Mg by Ca, Sr, or Ba led to expansion of (001) to 15.6-16.4A upon
glycolation. However, he noted exceptions in clay materials. Hathaway
(1955) describes a dioctahedral 14A clay mineral that does not expand
beyond 14-15A when Ca-saturated and treated with glycol. He feels that 43
.15 JO .05 0
CHLORITE
1 KAOLINITE
VERMICULITE
MICA
CA- MONTMORILLONITE
.30 .25 .20 .15 JO .05 0 sin 6 — 2i A 2.5% i 3.3 i A 5A i loX i
Figure 5.--Position and relative intensities of basal spacings of oriented clay-mineral aggregates. (After Bradley, 1954.) 44 this is more properly described as a vermiculite than a chlorite be cause the structure collapses at moderate temperatures. On the other hand, Grim and Johns (1954) describe a 14A clay mineral with the at tributes of vermiculite which expands to 17A upon glycolation. It is ap parent that no single set of properties characterize materials called vermiculite. It may be dioctahedral or trioctahedral, and it may or may not form two-layer organic complexes.
Walker (1957, 1961) suggests that failure of fully hydrated Mg- vermiculite to expand beyond 14.5A when treated with glycerol is char acteristic. According to him, the Mg-saturation plus glycerol treatment
is the only universally valid swelling test known at present.
The dehydration mechanism of vermiculite has been studied by
Walker (1956). The partially hydrated Mg lattice (14.36A), which is
stable over a wide range of temperature and humidity, is gradually con
tracted at higher temperatures to 9-10A. The temperature of irreversible
dehydration varies, and it may be greater than 500°C if an interlayer
cation of high hydration potential is present, e. g., magnesium. At
110°C an 11.6A phase appears which rehydrates almost instantaneously
in a humid atmosphere. This phase may be preserved if sealed in a
capillary or X-rayed in a controlled atmosphere immediately after heat
ing. Wars haw, Rosenberg, and Roy (I960) present useful data on changes
in layer silicates effected by heat treatment. 45
Three tests were used when vermiculite was suspected: (1) the
Mg plus glycerol saturation, (2) preservation of an 11.6A phase after heat treatment at 110°C, and (3) the relative intensities of 001 and 002 reflections (fig. 5).
Kaolin Minerals
Kaolinite
Kao Unite is the least complex clay mineral; however, its presence in mixtures may be difficult to evaluate by X-ray diffraction.
It is a very pure mineral, and published analyses deviate sUghtly from the theoretical values siUcazalumina = 2:1; SiOg = 46.54 percent;
AlgOg = 39. 50 percent; HgO = 13.96 percent (Grim, 1953). Small amounts of iron, alkalis, alkaline earths, and titanium are reported and seldom exceed a few percent of the total oxides. SoUd substitution is negligible, and it is beUeved that minor components may be present as separate phases, e. g., anatase, ilmenite, and leucoxene (Grim,
1953; Murray and Lyons, 1960). Kaolinite, halloysite, dickite, and nacrite are essentially polymorphs, although there is some difference
in degree of hydration. Nacrite and dickite are so rare in sediments that they are not considered here.
Kaolinite is a 1:1 layer siUcate with d(001) = 7 .15A. Well
crystalUzed material imparts an integral series of strong basal reflections 46 to high orders and also good general reflections. Disordered (% poorly crystallized, = randomly stacked) kaolinites are common (Brindley and
Robinson, 1946; Murray and Lyons, 1956; Brindley, 1961b). Disorder ing is manifest by hkl reflections with k 3n weak or missing. Such extinctions have been interpreted as, but are not limited to, translations of structural units along the b-axis by an amount equal to nb/3. Murray and Lyons (1956) present X-ray data for kaolinites ranging from well- crystallized to completely b-axis disordered material. The general trend is the extinction or merging of many weak reflections and the broadening of stronger lines. Disordered kaolinites are recognized by the lack of reinforced scattering in the range 3.5-2.5A where reflections
have the index k = 1 or 2 and by a strong 4 .4A(02) peak with a steep low-
angle termination (fig. 7). The I(001)A(021) *s roughly halved in dis
ordered kaolinites and a similar relationship is found in halloysite.
The coincidence of kaolinite 001 and chlorite 002 reflections at
7.15A requires auxiliary diffraction patterns to evaluate mixtures of
these minerals. Diagnostic tests usually involve differential solution
of chlorite in HCL or HgSO^ or thermal treatment with preferential re
flection shifts and intensity changes. Martin Vivaldi and Rodriguez
Gallego (1961) treated a 0.5 gm mixture of illite-chlorite-kaoUnite with
30 ml 20 percent sulfuric acid for 30 minutes and the resulting diffrac
tion pattern yielded no 14A peak, whereas iUite and kaolinite basal
peaks were intensified. Disadvantages of acid treatment are that other 47
»in 0
Figure 6. —Relative basal intensity curve of 2:1 dioctahedral clay minerals. Compiled from varied sets of montmorillonite- organic liquid complexes. (After Bradley, 1954.) .
R2;
Figure 7. --Effect of heating on kaolinite from Bath, South Carolina. Note complete loss of 001 reflection after treatment at 600°C for two hours. 48 clay minerals may be attacked, e. g., hectorite and well crystallized chlorites remain unaffected by concentrated acids (Ostrom, 1961). In addition, many important nonclay minerals in the clay fraction, such as zeolites, may be destroyed. Thermal treatment is non-destructive and provides a reliable means of evaluating the kaolinite (001) chlorite (002) superposition. Kaolinite dehydroxylates at about 550°C by a reaction of the type OH + OH = HgO + O and becomes amorphous to X-rays (fig. 7, metakaolin). Heating at 590°C enhances the chlorite 001 reflection and diminishes the 002 and 003 reflections through loss of interlayer scatter ing matter (Bradley, 1954). This one test usually suffices to establish the presence of small amounts of well-crystallized chlorite in mixture.
Bradley (1954) notes that after treatment at 400°C some chlorites lose intensity and the 7A feature is moved to a smaller angle and the 3.5A feature to a higher angle relative to kaolinite. This is interpreted as the collapse of poorly constituted brucite layers in chlorite and a mixing of 10A and 14A layers. The low angle feature is a combined 14A(001)/
10A(001) reflection. Therefore, the appearance of a low-angle bulge on the 7A reflection after heating indicates a superposition of chlorite and kaolinite in the original pattern.
Some investigators rely solely on the presence of an 003 chlorite peak at 4. 7A to prove its existence in mixtures. The presence of
analcime, with a strong reflection at 4.85A, precludes the use of the
third order chlorite reflection in this study. Even in the absence of 49 zeolite the chlorite 4. 7A peak appeared only as a shoulder on the high angle side of the strong 5A illite peak, and heat treatment was necessary to unequivocally establish its presence. Finally, interpretation is aided by the presence of a third order kaolinite peak at 37. 5° 28. Chlorite gives little or no intensity at its sixth order position, whereas the kaolinite 003 reflection is about one-eighth as intense as its 002 reflec tion (Bradley, 1954).
Halloysite
Halloysite may be represented by the structural formula
AlgSigOgtOH)^' nH20, where n ranges from 0.0-2.0. Naturally oc curring fully hydrated halloysite (endellite of some workers) contains one water layer between kaolinite structural units and exhibits a basal spacing of 10. LA. This may be expanded to 10.5-11.0A with organic molecules, or it may be irreversibly dehydrated to 7.2A at 110°C
(Brindley, 1961b). A metastable form with a basal spacing of 7.5-
7.6A is relatively common (Brindley, 1961b).
Fully hydrated halloysite is easily recognized by the 10.1A and
3.4A reflections. These peaks can be shifted by appropriate treatment.
The partially hydrated form is difficult to distinguish from kaolinite, but expansion normal to (001) may be accomplished by saturation with a polar liquid. The dehydrated form has almost the same diffraction 50 geometry as kaolinite and may be impossible to distinguish by X-ray methods in mixtures.
Halloysite is characterized by a high degree of stacking dis order along both a and b axes. Electron micrographs show that hydrated halloysite occurs as rolled tubes with the tube axis most commonly the b crystallographic axis and the unrolled tube the a-b plane. Basal re flections are weak because the tubular morphology prevents preferential orientation and because of the small number of layers forming the tube walls perpendicular to (001). Bates and Comer (1955) measured cylin ders 250-275A in diameter with walls only 25-50A thick.
Dehydrated halloysite mixed with kaolinite may be inferred by the shape of the basal reflections and the intensity of the (02, 114 ) band near 4.4A. In ordered kaolinites I(001)/l(020) is about two. Halloysite greatly increases scattering at about the 4.4A position relative to basal reflections, and when the intensity of the (02, 114) band equals or ex ceeds basal intensities, halloysite is probably present (Grim, 1953;
Molloy and Kerr, 1961; Brindley, 1961b). Mixtures of halloysite and kaolinite from Les Eyzies, France, give an X-ray pattern almost identical to pure kaolinite (Brindley and Comer, 1956). Electron micro graphs showed an abundance of tubes, and an artificial mixture of kaolinite and halloysite (1:1) gave almost the same pattern as the Les
Eyzies clay. In the presence of kaolinite, as much as 60 percent or even more of halloysite may be overlooked in an X-ray diagram 51
(Brindley et al., 1963). Conversely, electron microscopy has its problems, and it is possible to "lose" a fair proportion of kaolinite in a sample with a high percentage of tubular forms. It suffices to say, therefore, that both X-ray and electron microscope data should be in tegrated judiciously in kaolinite-halloysite determinations.
Chlorite
The chlorite structure is similar to that of biotite except that the 2:1 mica units are bonded by a brucite layer rather than a potassium ion. The resulting mineral is trioctahedral with talc sheets
(MggSi^OiotOHg) and brucite sheets (Mgg(OH)g) forming a regular a lte r nation of tetrahedral layers and octahedral layers in a ratio of 2:2. The generalized half-cell formula is (Rg^ Alx_z) (Si4 _x Alx)Oio (OH)g where R+2 _ Mg and Fe and R+^ = Fe. The negative charge in the tetrahedral position is compensated by R+^ and Al in octahedral posi tions or, where R+3 (VI) exceeds Al (IV), by deficiencies in octahedral positions (Foster, 1962). Saponite and vermiculite are similar to chlorite, but in the latter there is a complete brucite layer developed rather than a less densely populated hydrated magnesium interlayer
(Warshaw and Roy, 1961). Chlorites normally exhibit a stable 14A basal spacing because of the high layer charges, whereas chemically
similar montmorillonites and vermiculites are readily expanded along
[001] due to weak interlayer forces. The greater hardness of chlorites 52 over the micas is a result of a high opposite charge on the talc and brucite layer, plus the additional O-OH bonding mechanism between basal oxygens of tetrahedral sheets and the interlayer hydroxyl ions
(Brindley, 1961c). I For clay-mineral investigations the working parameters of chlorite are a monoclinic C cell, a = 5.33A, b = 9.23A, and d(001) =
14.2A (Brindley, 1961c). Disordered structures are known, and random displacements of structural units by nb/3 are indicated by re flection k V 3n weak or missing. Single crystal studies show polytypism analogous to that of micas, but recognition of stacking sequences de pends on reflections k ^ 3n not usually recorded on powder patterns
(Brindley, 1961c). Nelson and Roy (1954) have shown the existence of a whole series of chlorite polymorphs with 7A basal spacings that they
call septechlorites. However, these polymorphs bear greater strucr tural similarities to the kaolin minerals than to the chlorites, and the term has met some resistance. (See comments by Brindley in Nelson
and Roy, 1954; also Brindley, 1961c.)
Chlorite was identified in the playa and soil clays by an integral
series of basal reflections with a d(001) of about 14.2A (fig. 5). In un
treated flat-layer specimens the weak chlorite 001 was not resolved
from the long spacings of niontmorillonite, and diffraction from 7A
planes in chlorite could not be distinguished from the kaolinite 001 re
flection. All samples were heated to 590°C + 10°C for 2 hours and 53 air-quenched. This reduces kaolinite to an amorphous to X-rays con dition, irreversibly collapses montmorillonite, and partially dehydrates chlorite. The chlorite 14.2 peak is enhanced, and the second and third- order reflections are diminished. Thus, by selective heat treatment one can distinguish between basal peaks produced by 1:1 or 2:2 struc tural schemes. Anomalous reaction to heat treatment has been noted for amesite, chamosite, and chrysotile, all of which are relatively rare
(Nelson and Roy, 1954). In addition, thermal reaction temperatures may vary with grain-size, degree of crystallinity, and chemical com position. Grim and Johns (1954) describe a poorly constituted chlorite that yielded diminished 3 .5A and 7A reflections when heated to 450°C.
Well crystallized chloritic material is modified only at higher tempera tures best suited for the material being investigated. The chlorite con tent of the core was so slight that slides heated to 400°C gave patterns virtually indistinguishable from those heated to 590°C except for the
loss of kaolinite at the higher temperature. Soil samples gave best
chlorite peaks when heated to 590°C for two hours. The other char
acteristics which were used to corroborate the presence of chlorite
have been discussed. (See section on kaolinite.)
Inter stratified Minerals
Inter stratified clay minerals, commonly called mixed layers
(= mix layers), are mixtures of two or more different structural units 54 or species within a single crystallite. This occurs because the surface \ oxygen network of the different clay minerals is similar if not identical.
The problem facing the analyst is to determine the kinds of inter strati fied layers, their relative abundance, and the degree of ordering pres ent.
Thr6e kinds of ordering of mixed layers are recognized. The first consists of a regular alternation of layers with a definite repeat distance AB-AB-AB-, considered to be valid mineral species.
Rectorite, regularly alternating pyrophyllite-vermiculite layers, and
corrensite, regularly alternating vermiculite-chlorite layers are ex amples. Randomly mixed layers are those in which one layer has no
detectable influence on the neighboring layers (MacEwan, Ruiz Amil,
and Brown, 1961). In a mixture of equal proportions of A and B, the
chance of A or B occurring at a given position from an origin is equal
to its chance of occurring at all. Type three is the segregation of com
ponents into zones within a given crystallite and approaches the nature
of a mechanical mixture of minerals. Mixed layering is gradational
and may show a tendency toward regular alternation on the one hand or
segregation into zones on the other. Random mixing is by far the most
common phenomena and is the major concern of the ensuing discussion.
Mixed-layer minerals yield non-integral series of basal re
flections. These are. combined peaks and may not be indexed, but they
are commonly denoted by the indices of the component spacings producing 55 the reflection. Thus, the 001/001 peak of an illite-montmorillonite in terstratification may be the combined 10A/17A, 10A/15.4A, or 10A/12.4A peaks. Similarly, the 002/003 peak of a 10A/17A inter layering results from the combined 5A/5.7A reflections. The position of combined peaks varies almost linearly between closely spaced reflections. As the distance between extremes becomes greater, that is as the distance in reciprocal space increases and not d spacing, the curve becomes S- shaped (fig. 8). Peaks become diffuse in the steep part of the curve and may be absent if the middle portion approaches vertical (MacEwan,
Ruiz Amil, and Brown, 1961). The theoretical basis for scattering from mixed layers was first developed by Hendricks and Teller (1942).
The method involves the construction of theoretical scattering curves for assumed mixtures and matching observed diffraction patterns by trial and error. MacEwan, Ruiz Amil, and Brown (1961) and Brown and MacEwan (1949) present a number of these synthesized curves for various two component mixtures.
'A more straightforward method, which is adequate for most purposes, is that of Mering (1949). A line drawing scaled to l/d
(= 2 Sin 0/lambda) is marked at the positions where scattering maxima for the assumed pure components would appear (fig. 9). Observed dif fraction peaks are marked on the same scale and should occur at the same relative distance from the nearest reflections of the pure com ponents if assumptions about the interstratification are correct. The 56
00/001 003/005
001/002 0 0 3 / 0 0 6
0 0 2 / 0 0 4
------PROPORTION 10 A ------
A. Peok miration f o r I0A /I7.5A mixed layers (Moc Ewan, 1958). Dashed curve where peak not observed. The curve marked (W) is from Weaver (1956).
001/00'
0 0 2 / 0 0 3
------PROPORTION 10 A ------
B. Peak migration t o r I0A /I5.4 A mixed layers using Hendricks end Teller's formula
(Mac Ewan, et el.,1961).
Figure 8. —Curves of basal peak migrations for illite- montmorillonite mixed layers. Figure 9.—Mering's (1949) solution for random mixed- layer illite-montmorillonite.
Core sample from a depth of 36 feet, illitermontmoril- lonite = .25:. 75. Heavy lines are positions of peaks for pure components, dashed lines are positions of diffraction maxima for mixed layers, r* = l/d. I____ I J ______L 15.4 A/n 10
Calculated
15.4 A/n 10 A/n
Observed
17 A/n J ------L 10 A/n
Calculated
17 A/n 1 j ______L 10 A/n
O b s e rv e d
—i 8 10 14 16 26 28 30
r* (A ‘ x 100) m -3 58 peaks become broader and more diffuse as the distance between paired maxima increases. Closely spaced maxima, such as a second order
10A reflection and a third order 15.4A reflection, would give a sharp
002/003 reflection. The amount of each spacing present is proportional to the distance from the nearest maxima of each pure component pro ducing the peak. Figure 9 shows the calculated and observed values for a mixed-layer assemblage of 75 percent montmorillonite and 25 percent illite. Appropriate shifts and proportional distances were maintained when the expandable component was dilated to 17A with ethylene glycol or heated at 190°C for two hours. "It is not to be expected that the above process can be done exactly, but if it cannot be done at all the assumptions about the mixture are probably in error" (MacEwan, Ruiz
Amil, and Brown, 1961).
Examples of analyses and interpretations are given by Bradley
(1945, 1953), Weaver (1956), and for more complex three-component mixed-layering by Weaver (1956).
Semi-Quantitative Analysis
Quantitative evaluation of clay-mineral mixtures by X-ray diffraction is fraught with analytical difficulties. In its simplest form the diffracted intensity from a selected plane of a mineral is assumed to be proportional to the amount present. If pure standards are avail able one can prepare accurate calibration curves for binary mixtures 59 where one component acts as matrix for the other. In multi-component mixtures analysis is more complex, but with standards and the deter mination of mass absorption coefficients the method is feasible. The problem in clay mineralogy is largely one of selection and availability of standards, plus the burden of handling complex mixtures.
The variables of most serious consequence to the clay min eralogist are degree of crystallinity, chemical substitution, grain-size, and amorphous components in natural mixtures. Some of the problems presented by these variables may be obviated by determining relative clay-mineral abundances from comparison of basal peak intensities.
The method as outlined by Johns, Grim, and Bradley (1954) utilizes in tegrated intensities (on a count basis) of prominent basal peaks with a form-factor correction. Thus they estimate that the reflected intensity from an expanded component at 17A exceeds by a factor of four the scattering from similar material at 10A (fig. 6). The ratio montmoril- lonite-illite (M:I) is then Intensity (montmorillonite 17A)/ Intensity
(illite 10A) x 4. Similar.comparisons are made for kaolinite (002):illite
(003) and chlorite (004):illite (003) with a sm all correction for quartz contributing to the illite 003 reflection. According to Brindley (196Id):
"While this method is exceedingly crude . . . . it is reasonably sound
in its approach, and for exceedingly poorly crystalline materials, it is
probably the best that can be devised." Many variations of the foregoing
have been employed in clay-mineral research (Murray, 1954; Weaver, 60
1958a; Schultz, 1960; Freas, 1962; Griffin, 1962). The results of these subsequent studies point out that calculated form factors, i. e ., scattering ability of a particular mineral at a given angular position 26, may not agree with experimentally determined scattering factors because of
variable crystallinity and chemistry. Thus Freas (1962), when studying
the Eocene Golden Valley clay deposits, experimentally determined
working curves showing the kaolinite 7A peak and the montmorillonite
15.5A peak were three times the area of the illite 10A peak when the
minerals were present in equal amounts. Schultz (1960) compared poor
ly crystallized and well crystallized kaolins to illite and found the kaolinite
001/illite 001 varied from one to three for 1:1 mixtures. The author in
dependently found similar results with synthetic mixtures of A. P. I.
standard clay minerals.
The abundance of poorly crystallized and mixed-layer material
in the core and soils of the Willcox area compounded the usual difficulties
of semi-quantitative analysis. Montmorillonite-illite ratios were deter
mined after the method of Johns, Grim, and Bradley (1954) from the
diffractometer trace of the calcium-saturated and/or glycol-saturated
mounts. Inasmuch as the montmorillonite reflections covered an ap
preciable angular range, peaks were compared on an areal basis rather
than peak heights. This was done by constructing a smooth background
line with a french curve and measuring with a planimeter the area of the
basal peaks projected to background. The shape of the background 61
curve was determined by the diffractometer trace in the low 0 region when all minerals with spacings greater than 10A had been collapsed.
The area of the illite 001 reflection was determined after the low-angle
"tail" had been symmetrically related to the high-angle slope of the
peak. Bradley (1945, 1953) estim ates the scattering from m ontm oril-
lonite at 17A exceeds by a factor of four the scattering from illite at
10A. Coherent scattering in the low-angle range is subject to a rapid
decrease in the Lorentz and polarization factors and corrections need
be applied for this. Ratios are determined on well-crystallized
glycolated samples by
. peak area montmorillonite (17A) (X) peak area illite (10A) x 4
and for a montmorillonite peak centered at 15.2A the correction for de
creasing L. p is applied where with Cu radiation
peak area montmorillonite (15.2A) (2) peak area illite (10A) x 4 x ^•P jj* L .p 2.600
As a check against the values obtained, an alternative method was in
voked to determine M:I. Montmorillonite collapses to about 10A upon
heating at 200° for several hours (Grim, 1953, p. 98). The amount
present contributes proportionately to the increase in the 10A illite
peak; hence, measurement of the 10A peak before and after heating
gives a measure of the collapsible component present. The 10A peak 62 of the glycol-solvated slide should be measured and the same slide de hydrated at 200°C and irradiated under the same experimental condi tions with controlled humidity (Schultz, 1960). If the clay cake wrinkled when glycolated another slide was prepared or the peak areas of the calcium-saturated mounts were compared. In a few instances peaks of different preparations were compared using analcime or kaolinite as an internal standard to equate the difference in preferential orientation be tween slides. Generally speaking, this latter procedure is not very satisfactory.
The additional scattering at 10A from collapsed montmorillonite should not exceed one-fourth the area of the expanded 17A peak as this is the reduction in the form factor at that value of sin 0/lambda. Schultz
(1960) interprets an increase in the area of the 10A peak greater than
. 25x(area 17A peak) as due to collapsed mixed-layer minerals. In the core samples such an interpretation ejqjlains discrepancies in M:I as determined by the two methods. In core sample 106, for example, the
M:I determined by measuring the 17A and 10A peaks was 1:2, whereas the ratio determined by the increase in the 10A peak after heating was
1:1. Such a high proportion of inter layered material is readily seen on
the diffraction pattern as diffuse scattering between 10 and 17 angstrons.
Estimation of chlorite and kaolinite presents special problems
because of superposition of 00J2 reflections and mineralogical variability.
It was not feasible to determine working curves for the relative 63 scattering ability of the minerals encountered in this study because of the wide range of clay minerals analyzed, i. e ., soil, alluvial, and playa clays. Chlorite was estimated from the area of the (001) peak on the pattern of the sample heated to 600°C. The value thus obtained, relative to illite, probably represents a minimum because of the low structure amplitude of the peak (fig. 5). Kao Unite (001) and ilUte (001) were compared directly, and experience indicates that the value obtained represents a maximum since kaolinites with even a small degree of order have reflecting abiUties greater than equal amounts of ilUte in mixture. When kaoUnite and chlorite were mixed the 7A feature on the diffraction pattern represented scattering from both minerals. In this
case an estimated correction was appUed for chlorite addition to the
kaolinite first-order peak.
The ratios were then recast to show the proportion of each
clay mineral present in 10 parts of the total clay-mineral assemblage,
rounded to the nearest 0.5 part. This is the same procedure used by
Johns, Grim, and Bradley (1954) except that they arbitrarily aUowed
10 percent for poorly crystallized phases not accounted for in the dif
fraction pattern. Actual percentages are not inferred because nonclay
minerals commonly occur in the -2 micron fraction.
The author makes no pretense for high accuracy in the values
calculated, but the relative abundances should withstand scrutiny.
Pinsak and Murray (1960), in evaluating a similar analysis, state: "In 64 view of the many uncontrollable factors that must be considered, the quantitative estimates of individual clay mineral components can only be generalized. Assumptions have been made in the estimates and ac
curacy of the method is not great; but it is felt, in view of similar anal yses that have been made, that magnitude of error is not so great as the
amount of allowable error in estimation of relative abundance of clay
minerals in any specific sample."
Chemical Composition by X-Ray Diffraction
Some information concerning chemical composition of layer
silicates may be gained from data on b-axis periodicities and relative
intensities of basal reflections. The cation occupancy of the octahedral
layer determines the b dimension, and this figure is most reliably ob
tained from the (060) spacing. If the cell is monoclinic, b = d(060) x 6.
Trioctahedral clay m inerals have (060) spacings of about 1 .5 3 -1 .55A,
which is in the angular range 59.6°-60.6° 28 with CuKa radiation.
Dioctahedral minerals have a more contracted octahedral layer, and
d(060) lies between 1.49-1. 51A. Radoslovich's (1961) work on the cell
dimensions and symmetry of layer lattice silicates has firmly established
the dependency of the b periodicity on the nature of octahedral filling.
The (060) spacing is readily obtained from recordings of random powder
specimens. These were prepared by grinding the dried sample to about
200 mesh and packing the powder in an aluminum specimen holder. 65
Random orientation of platy minerals is more difficult to obtain than preferred orientation, and any shearing or smoothing of the powder in the holder should be avoided. For minimum orientation Brindley (1961a) recommends a specimen holder with a circular hole packed from the back with slight pressure. A neat scheme which claims excellent results in volves mixing the clay with therm oplastic cement (Lakeside No. 70C) and then grinding to size and packing in a holder (Brindley and Kurtossy,
1961). Quartz has a weak reflection at 1.54A that may cause errors if the presence of that mineral is forgotten.
The usefulness of basal reflections for characterizing micas was first noted by Nagelschmidt (1937) who showed the main difference between muscovite and biotite-phlogopite powder photographs was the
intensity of the 5A reflection. The relative intensities of basal reflec
tions can be used to distinguish octahedral cations with significantly
different atomic scattering factors. Brown (1955b) has calculated per
tinent partial sums of scattering amplitudes for micas with varying
chemical composition by the relationship
F(00£ ) = 22fa cos 2 itj£ z (3) ✓
where fa is the scattering factor of the atoms and z is the fractional co
ordinate of the atoms in a structure with a plane of symmetry parallel
to the layer. Brown treats the O-OH anion framework as a constant,
and Si, Al, and Mg are grouped together, as they have similar scattering 66 factors at low Bragg angles. In tetrahedral sites all atoms are treated
as Si, even though some A1 is present, and in the octahedral sites all
atoms are treated as Al, even though some Mg is present. Ferric iron
is treated separately as it is the only common cation affecting basal in
tensities in dioctahedral micas. The contribution from each part of the
structure to F(00 J2 ) is presented in Appendix A. If one wishes to find
F(002) for a mica of composition gHgOg g(AlgFei)(SiAl)g02Q(0H)^,
from Appendix A
/ F(5A) = F(5A) skeletal (0-0H) + F(5A)(Al3Fe1) + F(5A)
(Kj 5H2O0 g) r -120 4- 54 + 30 s -36
or in the case of total octahedral occupancy by Fe+^
F(5A) = -120 + 90 + 30 = 0
The latter case was chosen to illustrate the fact that as the Fe+^ in
octahedral sites goes up the structure amplitude for a 5A reflection ap
proaches zero. Relative intensities are obtained by
where >—< is the Lorentz and polarization factor. The relative in
tensities of basal reflections for dioctahedral micas of varying Fe+^ and
Al+^ content are presented graphically in figure 10. These have been
calculated for both the Debye-Scherrer and single crystal Lorentz and 67
S Single-Crystol
0 0 3
002 001
I'lFooJ'Z
003 001 — 002
4Fe 3Fe IA1 2Fe2Al Fe3Al 4AI Figure 10. —Change in intensities for dioctahedral micas. A. Lorentz-polarization factor for single crystal. B. Lorentz-polarization factor for Debye-Scherrer method. 68 polarization factors. Grim, Bradley, and Brown (1951, in Grim, 1953, p. 96) have plotted intensities of basal reflections using the Lorentz and
polarization factor for a single crystal. The decline in intensity received
from a slit system seems to follow powder laws more closely than those
for single crystals (Johns, Grim, and Bradley, 1954; Bradley, W. F.,
written communication, 1963), and therefore values for both have been
plotted (fig. 10).
One must be cautious in interpreting such data because of the
effect that mixed layering, orientation, and prolonged grinding may have
on basal intensities. Bradley and Grim (1961) observe, however, that
with clay-grade micas it is frequently possible to infer essentially
aluminian or essentially ferrian compositions. MINERALS OF THE WILLCOX PLATA
Core Material
Elite
The clay mineral composition of the core in terms of the prin cipal phases present is shown in figure 11. Elite is the dominant con stituent with varying but lesser amounts of montmoriEonite and vermiculite. Only traces of chlorite and kaolinite were detected. Dif- fractograms and photographic recordings of random powders yielded most intense 060 reflections in the dioctahedral range 1.51-1.52 ang stroms. Less intense and sometimes diffuse scattering was found in the trioctahedral range 1.53-1.54 angstroms. These values were ob tained from powders of the raw core material,, the -2 micron fraction, and photographs of the acid-treated sample. The intense scattering from a layer structure of dioctahedral composition is attributed to
Elite. Trioctahedral analogs of iHite are known, but they are rare and unstable in low-temperature mineral assemblages (Warshaw and
Roy, 1961).
Determination of the mica polytype from diffraction patterns of oriented aggregates proved unreEable because reflections for
69 70
Ktidinile
O Contains trees of chlorite
Montmorillonite Vermiculite CLAY MINERALS OF THE CO R E
Koolinite
O Contans trace of chlorite A Trace of chlorite a n d ,a koolinite
Vermiculite B CLAY MINERALS OF THE PLAYA SURFACE Figure 11. —Clay-mineral composition of the core and playa surface. 71 diagnosing successive layer rotations are of the type hh£ or Okl . In addition, calcite, feldspar, and analcime yield some interfering reflec tions in the critical angular range 20o-35° 26. Powder photographs did not resolve critical spacings any better than the diffractometer, and acid-treated samples yielded only the strongest lines usually recorded from illite. The poor film recording of acid-treated samples is perplex ing but may be due to absorption of X-rays by the amorphous residue of analcime not removed by washing, or to the partial breakdown of the illite structure by hydrochloric acid. The only consistent diagnostic spacing found was the (113) of 2M% muscovite. One would predict the presence of this polytype in the sediments of the playa because of the
large areas of silicic igneous rocks in the drainage basin.
mites of the core show no shift in peak positions but do exhibit
a distinct asymmetry of the 001 and 003 reflections so characteristic of
a small percentage of inter layered montmorillonite (Bradley, 1954;
Weaver, 1956). Basal peak heights were measured out to the 004 reflec
tion and normalized to intensity 001 « 100. With few exceptions, the
sequence is 001 >003 > 004 = 002, with 002 less than one-fourth the
height of the 001 reflection. In many samples the 002 reflections were
barely measurcable. These intensities agree with materials reported
as illite in the literature (Grim, 1953; Malloy and Kerr, 1961; table 1)
and imply a high content of ferric iron. 72
TABLE 1 X-RAY DIFFRACTOMETER SPACINGS AND RELATIVE INTENSITIES FOR ILLITE1./
ASTM 9-343 A. P. I. -35 A. P. I. -36 W. P.5 !5 (1) (2) (3) (4) dA *rel dA *rel dA *rel dA W 1M indices
10.0 100 10.0 100 10.0 100 9. 99 100 (001)
4.95 20 4.98 38 4.98 48 4.98 17 (002)
3.33 90 3.30 47 3.30 71 3.33 52 (003)
2.51 10 2. 54-2.49 27 2.54-2.60 40 2.49 14 (004)
1.986 10 1.99 28 1.99b 36 1.99 14 (005)
1. Carden Wood, A berdeenshire
2. Fithian, Illinois
3. Morris, Illinois
4. Core sample, depth 22 feet
V Diffractometer data from Malloy and Kerr, 1961 73
Montmorillonite
Montmorillonite in untreated oriented aggregates gives a broad
001 reflection from 12-13 angstroms. Calcium and magnesium satura tion produces a sharp first order peak at 14.2-14. 7A with a related rational sequence of higher orders (fig. 12). Another montmorillonite phase with a 15.2-15.4A basal spacing is present in most samples.
Both cationic forms of montmorillonite expand to 16.7-17.0A with ethylene glycol and to 17.6-18.0A with glycerol. Moderate heat col lapses the structure to 10A and with increasing temperature to 9.6A.
Experiments have shown that monoionic divalent montmorillonite complexes should imbibe two molecular layers of HgO and expand to
15.4A at humidities normally encountered in the laboratory. (See, for example, the isothermal curves of Mooney, Keenan, and Wood, 1952.)
Early workers assumed there to be a continuous variation in d(001) with water content, but it is now known that many reported spacings were produced by mixed hydrated phases (McAtee, 1956; MacEwan, 1961).
It was originally thought that the 14A Ca-saturated montmorillonite of the core was randomly inter layered Na- and Ca-montmorillonite result ing from incomplete filling of exchange sites by calcium ion. To test this, clay slurries were boiled in IN CaCl2 for six hours and excess
salt was removed by repeated centrifugation in distilled water. Sub
sequent irradiation showed no change in the 001 spacing of the 14A 30 28 26 24 22 20 18 16 14 12 10 8 6 4 2 14____12 JO 8 6 4 2 "28
Co + ethylene glycol
glycerol
z F igure 12. —Diffractometer traces of core sample 46, depth 36 feet. 75 phase. The same was true for the magnesium-saturated variety. Ran dom intercalations of 10A/15.4A layers in the right proportions (about
4:6) could produce the observed peak; however, dilation to full thickness in polar liquids virtually excludes this possibility.
A perusal of the literature revealed that ,,14A" montmorillonites are not unknown, particularly in young sediments. Nelson (1960) reports a 14.2A montmorillonite from bottom sediments of the Rappahannock
River that expands to 17. 5A with glycol. Gjems (1962) describes a curious soil-clay mineral that will not expand beyond 14A unless the slurry is first mixed with glycol and then glycerol added to the dried cake. This mineral has a high K4" fixing ability and has many properties of vermiculite. Bradley et al. (1958) show smoothed spectrometer traces of calcium and magnesium montmorillonites with two water layers and 15A basal spacings. Brindley (1955) discusses 14A ex pandable clay minerals, and Wars haw, Rosenberg, and Roy (1960, table
1) cite 14A montmor illonite and beidellite occurring in young sediments.
Two montmorillonite phases were found in sediments of the northern
Gulf of Mexico by Grim and Johns (1954). The 14.2A phase exhibits an
integral series of basal reflections and expands to 17A with glycol.
They believe that intercalation of brucite layers laterally at random on
an otherwise hydrated montmor illonite surface accounts for the observed
spacing. The brucite "islands" prevent complete collapse of the struc
ture at moderate temperatures but do not bond the layers sufficiently to 76 prevent expansion. The authors suggest that this phase represents chloritization of montmorillonite or degradation of chlorite. Mering
(1946) has shown that at about 30 percent relative humidity a skeletal double water layer forms in Ca-montmorillonites that gives a 14A 001 spacing. In this scheme, Ca+2 is octahedrally coordinated by water that forms a loose hexagonal network 4.4A thick between the layers.
At higher humidities a superimposed double water layer forms that is
5.9A thick. Similarly, Walker (1956, 1961) finds that the 14.3A stable
Mg-vermiculite contains an interlayer of loosely packed water molecules octahedrally coordinating magnesium. At high humidities the water forms complete hexagonal close-packed layers, and the basal spacing expands to 14.8 angstroms.
The observed diffraction phenomena indicate that the 14A montmorillonite of the core represents an incompletely filled inter layer region containing hydrated calcium and magnesium. The 15.4A phase observed in most of the recordings represents montmorillonite with a layer charge close to the ideal -.33 for each 0 ^q(0H)2 unit. The more abundant 14A phase has a higher charge in which separation of the individual layers is limited by the coulombic attraction of the hydrated cation and the negative silicate network. The attraction is not so great, however, as to prevent complete formation of glycol and glycerol com plexes. In terms of charge deficiency, the mineral lies somewhere be tween a low charge montmorillonite and a high charge vermiculite. 77
Cation exchange data on the pure phase would provide more evidence for proof, but unfortunately the mixture is too complex for purification.
An alternative explanation is to be found in the stepwise hydra tion process previously mentioned. At low humidities both Ca- and Mg- montmorillonites form one layer hydrates, and as the vapor pressure is increased two molecular layers are formed. The observed spacings could be ezqplained by interlayered 12.4A/15.4A hydrates in the propor tion 2:8. Mixed hydration states are considered very unlikely as most patterns were obtained at 45-60 percent relative humidity. In view of the many occurrences of 14A montmorillonites it is reasonable that ex pansible clay minerals possessing variable layer charges would be found in the young sediments of the playa.
Vermiculite
A moderately well-crystallized trioctahedral vermiculite oc curs in about half of the core samples. This phase is masked by re flections from other clay minerals and is resolved at 14.2-14.3A only after montmorillonite has been fully expanded with glycerol. It gives an integral series of basal reflections, and the intensity of the 004 re flection (7A) is very weak relative to the 002(14A) reflection. Detection of vermiculite was facilitated by the presence of reflections at 4 .7A(006) and 1.53-1. 54A(060). The structure collapses at 200°C with the loss of the sharp 14A peak and the appearance of an intermediate spacing. 78 usually about 10.8-11.0A. The latter spacing results from intermixing of collapsed and partially collapsed layers (Walker, 1961). Total col lapse occurs at 450°C, and the remaining structure cannot be resolved from that of illite. Some expansion of vermiculite was affected with ethylene glycol giving spacings of 14. 7-15.2A. Thus small amounts of vermiculite may be easily overlooked or interpreted as mixed layers, and it was found that the only consistently reliable technique for detec tion was the Mg plus glycerol test (fig. 12).
Chlorite
Nine samples contained a trace of poorly crystallized chlorite which appeared as a very weak 13.8A reflection on diffractograms of
slides heated to 600°C. These samples also gave weak 060 reflections
indicative of trioctahedral compositions. A few specimens heated to
200° gave weak 14A peaks combined with more intense 7A scattering.
This combination was attributed to chlorite as the relative intensities
indicate a form factor for a 14A mineral with high interlayer cation
content (fig. 5).
Kaolinite
Traces of kaolinite were recognized in only four samples near
the top of the core. The basal peaks were so weak that higher orders
were not resolved from peaks of other minerals or background. The 79 coincidence of chlorite and vermiculite peaks in the 7A angular range further complicated identification. The samples taken in the top eight feet of the core contained very little expandable material and gave d(060) spacings of 1.49-1. 51A attributable to kaolinite. Some weak 7A peaks remaining after moderate heat treatment could be assigned to kaolinite or chlorite. When diffraction data were ambiguous, chlorite(?) was chosen because kaolinite is unstable in alkaline environments.
Mixed-layer minerals
Inter stratification of expandable and non- expandable units was the rule throughout the length of the core. Such mixed-layering was manifest on diffraction patterns in two ways: (1) as diffuse scattering on the high angle side of the first montmorillonite peak, and (2) as a discrete but weak peak between the first order illite and montmorillonite reflections. Interstratification of the first type is not amenable to anal ysis and represents scattering from crystallites which contain adjacent layers of unequal thickness randomly mixed over a wide range of pro portions. This type was particularly abundant in the top 20 feet of core and produces asymmetry of the first-order montmorillonite peak easily seen as a tail or hump on the high-angle side. Mixed-layering of the second type was rare, but in several samples analysis by Mering's
(1949) method was possible. Fewer combined reflections were present than was desired, but good agreement was obtained between calculated 80 and observed spacings out to 003/004 for an illite-montmorillonite in terstratification of . 25:. 75 (fig. 9).
Analcime
Analcime is validly considered with the zeolite group of min erals, although its chemistry, structure, and paragenesis are closely related to the feldspathoids (Deer, Howie, and Zussman, 1963). It is cubic with a = 13. 7A and IGfNaAlSigO^^O per unit cell. Linked
(Si, Al)-0 tetrahedra are so disposed that large channels accommo dating water, and smaller channels occupied by Na, run continuously , O through the structure. Na is ionically bonded by the compensation A1 proxy Si+^, whereas water is loosely held in large channels by hydrogen bonding. It is a ubiquitous constituent of the playa sediments and is re markably consistent in its diffraction characteristics and abundance throughout the core. Initial identification was made from strong
211(5.57A) and 400(3.42A) reflections. Comparison with analcime from Nova Scotia shows excellent agreement in both peak position and relative intensity (fig. 13). An average of six powder photographs gave a . 13.67A for the playa material.
Examination of the sand and silt fractions with the petrographic
microscope disclosed no analcime in those size grades. Electron micro
graphs show analcime as equant cubic crystals in the -2 micron fraction
(pi. 2). These crystals are morphologically similar to those in electron 81
WP54
WPI06
ASTM
Figure 13. —Diffractograms of analcime from Willcox Playa and pure analcime from Nova Scotia. (Copper Ka radiation.) 82 micrographs of Noll (1944) and comprise about 10 percent of the samples photographed.
Before obtaining electron micrographs several samples were analyzed for analcime content by the method of "known additions"
(Brindley, 196Id). A small amount of the pure mineral to be determined is added to the mineral assemblage, and the reflected intensity before and after the addition is measured. If Wp is the weight proportion of component P before addition, Wj. the amount of pure component P added to the mixture, and and Ig the reflected intensities before and after the addition, then
wt • h Wp = ______% (5 ) Ii l +Wt-± h
If the ratio of average mass absorption coefficients before and after the addition is unity (U*]/U*2 = 1 ), then one can calculate Wp from equa tion (5) without analysis for U*. However, if the addition changes U*, then Wp will be in error accordingly. The difficulties in determining mass absorption coefficients may be obviated by calculating Wp for several values of Wt and extrapolating to Wt = 0 . Wp was calculated for several additions, 0. Igm/gm, 0.2 gm/gm, and 0.3 gm/gm, and the 83 values thus obtained extrapolated to Wj. - 0 for a more accurate per centage value.
To one gram of lightly ground sample was added tenth-gram increments of powdered analcime from Nova Scotia. The mixture was then ground three times in an agate mortar under acetone and washed into a 5-dram vial. A few grains of dispersing agent were added and five drops of suspension applied to the center of a 27 x 46 mm glass slide. After the slurry had dried, the flat-layer cake was masked so that the sample area intercepting the X-ray beam was the same on all slides. The slides were centered on the rotation axis of the goniometer and oscillated through the angular range of the (211) peak. Ten to fifteen peak heights were measured and averaged to determine Ij /L j. Figure
14 shows the results and extrapolations of these calculations. They are in good agreement with the visual estimates from the electron micrographs and give a range of five- to ten percent for analcime con tent. Inasmuch as the analcime peak intensities varied little from sample to sample, this figure also represents the magnitude of zeolite content for the entire core. The results also indicate that the addition of analcime to the natural mixture does not greatly alter the absorption of the sample. The mass absorption coefficient is 32.8 (water not in cluded), compared to montmorillonite (32), metakaolin (33), illite (51), and muscovite (43). 84
X Andclme W P-3 A
% Anolcim# WP-216 B
Wf (ym/gm)
% Anoldme WP-216 C
Figure 14. —Analcime content of two core samples by the method of known additions. 85
Calcite
Calcite occurs in all size grades in the core and constitutes as much as 80-90 percent of the sand fraction in some samples. It is pri marily, if not entirely, the product of authigenic crystallization and biological activity. Angular fragments of translucent sculptured ostra- cod valves are prominent in both the silt and sand fractions. Shell frag ments have served as nuclei for radial growth of calcite crystals with prismatic and scalenohedral habit. These crystals are very fine grained and contain much included organic matter. They exhibit parallel extinc tion, high birefringence, and effervesce vigorously in dilute hydrochloric acid. Aggregates of tiny crystals occur in several forms. White, sac- charoidal, spherical grains in the coarse fraction were found to be com posed of radial groups of calcite crystals. Much more common are light-brown^ spherical and rod-shaped aggregates of crystalline calcite, ostracod fragments, and detrital minerals. These range in size from
. 04-. 15 mm in diameter and show varying degrees of effervescence in acid. Some effervesce vigorously and appear to be aggregates cemented by calcite. Others effervesce sluggishly, and the nature of the insoluble
residue suggests these are fine mineral grains agglutinated by organic
matter. The rod-shaped pellets resemble closely the fecal secretions
reported in modern marine sediments (Moore, 1939). Several samples
contained reddish-brown spongy aggregates that had evidently been 86 oxidized and redeposited from dessicated margins of the lake. Overall, cemented and agglutinated grains formed the bulk of the sand fraction of the core.
Beautifully developed calcite euhedra appear at a depth of 80 feet and occur in varying amounts to the bottom of the core. They exist as elongate near-perfect single crystals with well-developed prisms and scalenohedrons terminated by rhombohedrons. Less commonly, rhombohedral crystallization occurs alone. Many are without inclu sions and perfectly clear, whereas others contain considerable amounts of organic matter. These crystals are definitely authigenic as they show little evidence of mechanical wear, and they completely outsize the detrital grains with which they occur. Crystals a few tenths of a millirpeter long were the rule, although some measured 1 - 2 mm long in sample 139 at a depth of 100 feet.
Quartz
Angular- to subangular quartz grains were identified in the coarse fractions. They are generally smaller than the aggregate grains, ostracod valves, and calcite crystals and show little evidence of long transport or reworking. Quartz is a minor constituent in the -2 micron fraction and appears as fractured fragments in electron micrographs
(pis. 1 and 2 ). 87
Feldspar
Analysis of the principal feldspar peaks on diffractograms failed to yield any prevalent distinction between members of this group.
Authigenic albite was anticipated on geochemical grounds, but it was not found. Plagioclase feldspar was notably rare in the coarse frac tions, and the writer identified surprisingly few grains in over a hun dred mounted and loose-sediment samples examined. Potassic feldspars were more common, and many grains exhibited the clarity and lack of twinning typical of sanidine.
Heavy minerals
Heavy mineral suites were separated from the sand and coarse silt fractions of the core and examined with a petrographic microscope.
Pyroxene and amphiboles were grouped (pyroboles), as were epidote, clinozoisite, and zoisita (E-C-Z). The frequency of mineral groups and species was visually estimated; the results are tabulated in Ap pendix B.
Primary bentonitic material overlooked in the visual examina tion of the core might be detected in heavy mineral suites. Euhedral apatite, zircon, and biotite, with or without sphene, are typical of silicic volcanic ash (Weaver, 1963). A few samples had so little coarse material that the heavy-mineral separates were too fine grained for 88 microscopie examination. The remainder contained abundant minerals with a decidedly mafic character. Zircon, tourmaline, rutile, sphene, and garnet are rare or present in small amounts; whereas pyroboles,
chlorite, biotite and E-C-Z are abundant. Tourmaline is the brown variety typical of low-rank metamorphic rocks, and strongly pleochroic brown basaltic hornblende forms a significant proportion of the total pyroboles. Dark-brown and green biotite are common, and much of
the brown biotite has been leached to a pale golden brown through loss
of iron (Folk, 1961). Sedimentary chlorite was distinguished from
green biotite by the mottled appearance imparted to chlorite from abun
dant inclusions. Except for a few zircon euhedra, all the minerals
showed the effects of weathering and mechanical abrasion.
A striking feature of the heavy-mineral assemblage is the
paucity of stable minerals and an abundance of more labile constituents.
The predominant minerals are those one would expect from a mafic-
igneous and metamorphic rock terrain. However, the low frequencies
of zircon, rutile, sphene, and garnet do not adequately represent con
tributions from silicic igneous rock areas. These minerals are the
most dense heavy minerals, and because of hydraulic factors operating
in this low-energy environment, they are not found in the coarse-size
fractions examined. It is felt that the frequencies of apatite better
represent the sediment contributions from silicic igneous terrains
which lie east and west of the playa. Even allowing for hydraulic 89 factors, it is apparent that the heavy-mineral suites are influenced mainly by detritus from mafic-igneous and metamorphic rock terrain to the north.
Mineraloid
A black, finely divided substance was associated with the min erals of the core. It oxidized rapidly, usually with a few minutes, with a color change from jet black to light gray. Treatment with hydrochloric acid generated a strong odor of hydrogen sulfide. This is the foul black
"stinking clay" described by drillers in the area. It is said by them to turn yellow when exposed to the air (Meinzer and Kelton, 1913).
Cuts from the sealed core were X-rayed before a visible color change took place. The diffraction pattern of the raw sample before oxidation showed no peaks other than those accounted for in the - 2 micron fraction. The negative Eh values, evolution of H2 8 , and oxida tion of sulfide(?) to sulfur (?) suggest this black material to be a ferrous sulfide, possibly hydrotroilite. It is a finely divided colloidal material found in many Recent muds and is believed to be of the composition
FeS- H2 O or FeS* nE^jO. Em ery and Rittenberg (1952) describe sim ilar acid-soluble iron sulfides in cores from harbors, estuaries, and bays along the southern California coast. 90
Electron Micrographs
Method
A small portion of the dried clay was ground in an agate mortar and pestle under absolute ethyl alcohol until dry. After the third grind
ing a 300 mg split was washed into a screw-top jar and 100 cc of dis tilled water added. The suspension was then placed in a Blackstone
Model SG-3 ultrasonic vibrator. The sample was subjected to vibra
tions in the 20 kc range for 15 minutes with the water in the tank and
jar at resonance level. Two drops of the dilute suspension were evao-
rated on a freshly split muscovite substrate and stored in a petri dish.
The mica substrate was masked and a platinum-shadowed
carbon replica obtained by evaporating platinum and carbon in vacuum
at angles of 30° and 90°, respectively, to the plane of the substrate.
The platinum-carbon replica was removed from the substrate, and any
clay adhering to the metal was dissolved by floating in hydrofluoric
acid. Electron micrographs of these shadow-cast replicas were made
for the writer by Professor R. L. Sloane of the University of Arizona
on a Philips EM-100B electron microscope.
Discussion
Plate 1-A shows a mechanically separated portion of the clay
fraction of the core from a depth of 32 inches. The larger angular PLATE 1
ELECTRON MICROGRAPHS OF SAMPLE 208, DEPTH THREE FEET
A
Mechanically concentrated analcime phase of the -2 micron fraction. Note the platy clay minerals in the lower right-hand corner of the micrograph; x 8,700.
B
Unknown fibrous mineral; x 11,000. 91
B 92 grains about 1 micron across are probably detrital quartz. The finer,
more equant material is analcime. Some of the very fine grains in the right-center of the photograph are also analcime. A few platy clay
grains appear in the lower right-hand corner. They are of special in terest because the edges have the embayed appearance of resorbed min
erals as seen in thin sections. They may represent partially dissolved
kaoUnite grains, as this sample contained the highest proportion of this
mineral, and kaolinite would be the least stable of the clay minerals in
the playa sediments. The percentages of the minerals in the micrograph
are not significant owing to the selective nature of the sample.
A large analcime crystal is shown in plage 2-B. None of the
crystals observed exhibit trapezohedral forms reported from other
occurrences (Milton and Eugster, 1959). Montmorillonite is seen as
fluffy masses and aggregates of tiny particles. Its morphology is af
fected by grinding, nature of the dispersing agent, and cation satura
tion. The appearance in plate 2-B is typical of H- and Na-saturated
montmorillonites (Bates, 1954).
lUite appears as thin crystals with edge intersection of 120°
in plate 2%A. Also present are rhombic grains, probably calcite, and
detrital quartz. The elongate crystal in the center of the photograph is
a synthetic phase of unknown composition. Note its growth around an
analcime cube along the lower edge of the crystal. PLATE 2
ELECTRON MICROGRAPHS OF CORE SAMPLE 125, DEPTH 90 FEET
A
Elite plates, analcime, and fluffy looking aggregates of montmorillonite; x 8, 700.
B
Large analcime crystal with montmorillonite. Note large Elite crystal in upper right-hand corner; x 7,000. 93
B 94
Unknown m ineral
A mineral with tubular or fibrous morphology is shown in plate
1-B. It appears to have the tube-in-tube morphology described by Bates,
Hildebrand, and Swineford (1950) in their definitive paper on halloysite.
The mineral is about ten microns long, and one end appears to be un rolling to give a spatulate form also known in halloysite. The mineral was tentatively identified as halloysite(?) by the writer. Dr. T. B.
Bates, of the Pennsylvania State University, kindly consented to ex amine the electron micrograph and offered the following observations:
(1) there is no real evidence for tubular morphology because the density of the edges and center of the fiber are the same, (2) the length/width ratio is greater than in most halloysite tubes, (3) the very regular, smooth edges are not found in most halloysite tubes where outer layers tend to peel off giving a more fuzzy, irregular outline. Some of this would be expected in a large crystal, particularly at the curve of the fiber.
Dr. Bates feels there is not sufficient evidence to call the fiber
halloysite. The length/width ratio was the principal reason that the
initial identification was queried, and this together with the discrepancies
noted by Dr. Bates are sufficient to negate the identification. Professor
R. L. Sloane (personal communication) suggested the mineral might be
either palygorskite or chrysotile. The same difficulties arise here in 95 that tubular morphology is common to both these minerals. On physio- chemical grounds, the environment of Lake Cochise and the surrounding terrain would favor the formation of palygorsMte - sepiolite type min erals, Unless better samples are found, the fibrous object in question is best relegated to the limbo of ’’unknown minerals. ”
Playa Surface
Elite dominates the clay fraction, as in the core, almost to the exclusion of other minerals except analcime (fig. 11). Expansible phases constitute less than five percent of the clay minerals. They are in order of abundance, mixed-layers, montmorillonite, and vermiculite.
A trace of kaolinite was found in one sample near the center of the playa and was queried in two others. Chlorite occurs in four samples and very probably exists as a minor phase in two more.
Also, analcime is an ubiquitous phase and is similar in diffrac tion characteristics and amount to the analcime at depth. Feldspar, quartz, and calcite occur in varying amounts without any significant variation from these same minerals in the core. Halite, thenardite, and burkeite are found in the soluble white crusts of the playa surface.
These minerals have crystallized from evaporating surface water and capillary ground water. 96
Soils and Alluvial Samples
General
Forty-two stream channel and soil samples were obtained within the drainage basin. Sample sites were selected in an effort to represent fine-sediment production from diverse geologic terrane and to ascertain the minerals in regolith favorably located for transport to the playa.
Mineralogy
The clay mineralogy of the playa surface and surrounding terrain is shown in figure 15. Samples taken in the extreme northern and eastern parts of the drainage basin have been located along the border of the map in order to preserve a convenient scale for the diagram.
Several difficulties were encountered in interpreting diffraction patterns, the most serious of which was the differentiation of chlorite and kaolinite.
Resolution of chlorite and kaolinite peaks is lost at low concentrations, and a standard procedure was adopted in such cases. When a strong 7A peak was present, the pattern was examined for the presence of a chlo
rite 003 reflection. If present, even though no 14.2A peak appeared upon heating, the 7A scattering maximum was assigned to both chlorite
and kaolinite. If a 14.2A reflection appeared upon heat treatment, all
the scattering related to 7A was attributed to chlorite in the absence of © © © © e © © ©
LEGEND
figure 15.—Clay mineral distribution in soils, stream channels, and the playa surface. 98 a kaolinite 003 reflection. If both third order diagnostic reflections were missing, a 7A peak was attributed to kaolinite-chlorite undiffer entiated.
Elite, montmorillonite, and vermiculite presented no special difficulties. Halloysite was suspected in several samples but not posi tively identified. Of particular interest is the presence of analcime in three soE samples. Samples 15 and 19, from the northeastern margin of the playa, contained appreciable amounts of this mineral. Sample 19 was taken from a borrow pit in which soil and two generations of older very coarse-grained alluvium were exposed. Analcime was found at all three levels. Cooper*s geologic map (1960) shows that these sample sites lie on or within the beach ridge marking a high stand of Lake
Cochise. This strongly suggests that the zeolite was introduced to these sediments when they were submerged during pluvial times. A similar explanation is invoked for analcime in sample 7 at the south end of the playa. If this mineral were detrital, one would anticipate its occurrence in other samples of this major drainage into Lake Cochise.
In a general way clay-mineral production and geologic terrane
can be associated. Elite was produced from all terranes and is dominant
in most samples, just as it is in the playa. Soils and channel samples
high in chlorite and vermiculite come from the Dragoon Mountains where phyllites, schists, and granite crop out. Northward, on the west
side of the playa, montmorillonite increases as the influence of the 99 granitic terrane of the Little Dragoon Mountains is felt. A soil sample from Texas Canyon in the Little Dragoon Mountains contained mont- morillonite as the dominant clay mineral. From the complex terrane north of the playa chlorite, chlorite-kaolinite, and montmorillonite are produced. Most of the chlorite-kaolinite samples are probably chlorite, because they come from gneissic and basic igneous rocks. These samples were all from stream channels and are a good representation of material being carried toward the playa from the north. The samples from Pre- cambrian granites in the Dos Cabezas Mountains (20, 21, and 23) are high in montmorillonite with minor chlorite and kaolinite. Sample 23 was taken from a thin residual soil on granite and contains more than
50 percent montmorillonite. Northward the mixing of sedimentary and extrusive igneous rocks is reflected in less montmorillonite and more chlorite and chlorite-kaolinite. South of the playa samples with relative ly high proportions of kaolinite and kaolinite-chlorite were found. Some of these samples may represent reworked lake sediments.
The lack of correlation between the clay minerals of the playa surface and the surrounding terrain is striking. The difference is still marked if only the stream channel samples and the playa surface are compared. Droste (1959, 1961a, b, c) determined clay minerals of the surface of 45 California playas and streams draining into their basins.
He concluded that there was no difference among the clay minerals of the playa, whether saline, brackish, or essentially fresh, and the clay 100 minerals supplied by the local drainages. Several factors operating in the Willcox Basin explain the discrepancy between Droste's findings and the present work.
(1) Many of the streams draining mountains to the east and west of the playa disappear on the bajadas and show no evidence of having re cently supplied detritus to the dry lake. Pine (1963) found no channels on the east side carrying debris to the playa; however, he did find sev eral buried channels with heavy minerals characteristic of Dos Cabezas terrain .
(2) Several feet of playa surface have been removed by wind action, and it is not unlikely that some clay minerals, particularly very fine colloidal montmorillonite, were selectively deflated.
(3) Leaching and oxidation of the playa sediments to a depth of at least eleven feet are indicated. Such leaching would selectively trans port fine colloids downward and strip potassium from illite. This would explain the low montmor illonite content of the surface and the relatively high mixed-layer content of the surface and near-surface materials.
The clay mineral distribution at the surface of the playa cannot be explained by normal sedimentary processes. Their uniformity sug gests that some form of quasi-equilibrium has been reached between geologic processes, mainly wind and meteoric waters, and the surface sediments 101
All of the clay minerals found in the drainage basin are found in the core with the exception of kaolinite. This problem is discussed in another section of this paper. ORIGIN OF PLATA SEDIMENTS
Detrital Minerals
Field and laboratory evidence favor a detrital origin for the clay minerals in the playa sediments. The absence of volcanic glass and heavy minerals associated with bentonitic montmorillonite pre cludes this type of neo-formation in Lake Cochise. According to
Weaver (1958b), the ability of expansible clay minerals to "fix" potas sium is of genetic significance. Thus, montmorillonite derived from high charge micas should take up K* and collapse to 10A, whereas this mineral aggraded from volcanic material would have a low charge and only partially collapse. Similarly, two inheritance groups in vermiculite are recognized: (a) a biotite vermiculite that collapses to 10.3A, and
(b) a non-biotite vermiculite that collapses to 11.0-11.7A with K* in the exchange positions. Whether these interpretations are always valid is debatable; however, K4" fixation is a measure of layer charge and can be usefully employed. Samples were immersed in IN KOH for 4 hours at 80°C and then washed in distilled water and methanol. Those samples in which 14A montmorillonite was the only expandable com ponent showed almost complete collapse to 10A with a small proportion
102 103 of mixed layering. Clays containing more than one montmorillonite phase, vermiculite, and chlorite showed appreciable mixed-layering after potassium saturation. Rationalization of these data indicate that there is a spectrum of layer charges, from high-charge montmorillonite and vermiculite to low-charge montmorillonite. The differences are ascribed to different degrees of weathering and weathering susceptibility of parental materials.
Vermiculite was derived from comminution of leached biotite as evidenced by abundant biotite in the heavy mineral fraction and by weak trioctahedral reflections on diffractograms. Chlorite is known to form from degraded illite or montmorillonite by absorption of Mg in the marine environment (Grim and Johns, 1954; Powers, 1954). Such chlorite is poorly crystallized and has a low thermal stability. Chlorite in the core has a high thermal stability and gives enhanced 001 reflec tions after treatment at 600°C. The optical and X-ray data indicate that the chlorite in the core is of igneous origin. In the source area chlorite is more abundant than vermiculite, whereas in the core the re verse is true. Leaching of chlorite to form vermiculite must have oc curred in transport to the playa.
Illite is the dominant mineral in the core and in the regolith of the source area. In the presence of potassium ion illite would form from stripped illites and mixed-layer illite-montmorillonite from beidellitic parts of montmorillonite (tetrahedral substitution). This 104 process occurs easily and fast, as is exhibited by soils treated with potash and Recent marine sediments. Lake Cochise was most probably low in potassium, and clay minerals aided in its geochemical disposaL
Nevertheless, the bulk of the illite in the core is detrital in origin.
Data obtained from optical and electron microscopy demon strate that quartz and feldspar are also detrital. The geochemical im plications of these data will be discussed later.
Authigenic Minerals
Analcime
This mineral occurs admixed throughout the playa sediments as extremely small equant crystals (pis. 1 and 2). Diagenetic or syn- genetic zeolites in sedimentary rocks form mostly in closed systems associated with volcanic ash. Of forty-three occurrences (Deffeyes,
1959), twenty five are associated with volcanic ash, and in several others a volcanic association is implied. To my knowledge, all pre viously reported occurrences in the western United States save two
(Keller, 1952; Hay and Moiola, 1963) are attributed to crystallization from devitrified volcanic glass. Closed systems include saline (playa) lakes, isolated marginal marine basins and lagoons, marine basins on shallow platforms, and deep-sea trenches. Zeolitization requires the ready availability of silicon, aluminum, alkalis, and alkaline earths, 105 particularly sodium and potassium. Zeolites formed at low tempera tures and pressures are clinoptilolite (= high silica heulandite), heulandite, analcime, phillipsite, laumontite, and mordenite (Deffeyes,
1959). Of these clinopilolite-heulandite and analcime are most frequent ly reported from sedimentary rocks. Mordenite is suspect as a dia- genetic product (Deffeyes, 1959; Coombs et al., 1959) and may be of hydrothermal origin as in tuffaceous sandstones and shales of northern
Japan (Sudo, 1959). Chemical formulae and pertinent ratios are p re sented in table 2.
Natural occurrences.—Authigenic zeolites have been reported from sedimentary accumulations from Devonian (Bur'yanova, 1960) to
Recent (Murray and Renard, 1891). A review of the literature showed that the basic mechanisms for low temperature formation of zeolites is limited to a few reactions. These will be summarized with pertinent references to such occurrences.
There are many reports of field evidence for corrosion of vol canic glass by alkaline solutions to form zeolites. The following evi dence has been offered for such a chemical mechanism:
(1) Pseudomorphic filling of glass shard molds by zeolite
(phillipsite, Hay and Moiola, 1963; clinoptilolite, Hay, 1963).
(2) Progressive replacement of glass or pumice by zeolite
(analcime, Bradley, 1929; clinoptilolite, Bramlette and Posnjak, 1933; TABLE 2 / COMMON AUTHIGENIC ZEOLITE MINERALS—
______Number of ions per unit cell______'
M ineral Si A1 Mg Na Ca K HgO Si/Al
Heulandite(l) 26.35 9.63 — — 2.63 3.54 — — 23.66 2.73
Analcite(2) 1.995 1.03 .006 .791 .061 .076 2. 014^/ 1.93
Phillipsite(3) 9.91 6.15 — — 1.93 0. 98 2.02 15.38 1.61
Erionite (4) 27.42 8. 78 .79 1.34 1.49 1.95 28.01 3.12
Mordenite(S) 19.61 4.31 .16 2.22 .96 .19 13.71 4.55
Laumontite(6) 15. 65 7. 70 — — — — 4.58 — — 12.96 2.03
(1) (Ca, Nag) (AlgSiyOig)' GHgO Amygdaloidal basalt, Lanakai Hills, Hawaii, based on 72 oxygens per cell.
(2) NafAlSigOg)* HgO Cyclopean Island, Sicily, based on 7 oxygens per cell.
(3) (l/2Ca, Na, K)g (AlgSigOig)" GHgO Cavities in basalt, Moiliili quarry, Hawaii, based on 32 oxygens per cell. .
(4) (Nag, Kg, Ca,Mg)4 o 5 (Al9Si27072)'2 7 H2 O Fibrous erionite, Baker County, Oregon, based on 72 oxygens per cell. 106 TABLE 2—Continued
(5) (Na2 , K2 , Ca) (Al2Siio024)**7H20 Amygdales in vesicular andesite, Idaho, based on 48 oxygens per cell.
(6) (Ca(Al2Si40i2)°4-3.5 H2 O In veins, Montana, based on 48 oxygens per cell, a/ Data from Deer, Howie, and Zussman, 1963. b / (OH) 107 108 laumontite, Ross, 1958; clinoptilolite, Deffeyes, 1959; clinoptilolite,
Hay, 1963).
(3) Replacement of tuffaceous beds by zeolite (analcime,
Bradley, 1929; erionite and heulandite, Regnier, 1960).
(4) The intimate association of volcanic ash with replacements, cavity fillings, layers, cementing, or admixed zeolite (analcime and phillipsite, Murray and Renard, 1891; analcime, Ross, 1928, 1941; heulandite, Gilbert and McAndrews, 1948; phillipsite, Zen, 1959; many zeolites, Bur'yanova, 1959; phillipsite, erionite, clinoptilolite, Deffeyes
1959; analcime and heulandite, Benda, Erd, and Smith, 1960; clinoptil olite and analcime, Coombs, 1960; phillipsite, Hay and Moiola, 1963).
Authigenic zeolite not associated with volcanic material is relatively rare. Proposed sources of silica and alumina for zeolitiza- tion are dissolution of quartz, plagioclase, and pyroxene and true or colloidal solutions. Keller (1952) reports a zone of analcime in the upper Chinle Formation of Utah that he believes is correlative with a similar zone in the Popo Agie Member of the Chugwater Formation in
Wyoming. The analcime is found with limestone and ocherous fine grained laucustrine sediments devoid of vitric material. He states that
Na-rich waters acting on a subadjacent bed of non-bentonitic montmoril- lonite provided the necessary cations. Analcime, with calcite, dolo mite, kaolinite, brammalite, and melanterite, occurs in a concretion from a coal seam at the base of the Monongahela Series in Pennsylvania 109
(Foster and Feicht, 1949). It is proposed that the Si and A1 for analcime came from downward percolating, highly alkaline waters acting on kaolinite. Coombs (1960) found zeolites formed from acid glass and detrital plagioclase on a regional scale in polymictic sediments from
New Zealand. Vernet (1961) reports a montmorillonite-analcime as semblage near Stanleyville, Congo. He states that although volcanic ash is not known in these rocks, other occurrences favor devitrification of glass to form analcime. Ross (1928) suggested that analcime in tuff- aceous sandstones near Wikieup, Arizona, could have crystallized from the dissolution products of clay minerals in an alkaline lake. He later
(1941) concluded that enough volcanic m aterial was present to provide the necessary constituents. Hay and Moiola (1963) found that felsic glass could yield the necessary cations for the formation of phillipsite in tuffaceous beds at Searles Lake, California. However, 21 of 72 samples contained authigenic analcime, and 36 of 72 samples contained authigenic K-feldspar, most of which is not associated with tuff beds.
Quartz, plagioclase, and pyroxene showing evidence of solution could have provided only one-tenth of the necessary silica and alumina. They suggest that the remainder was supplied by diagenetic destruction of montmorillonite.
The precipitation of zeolites from true and colloidal solutions is cited by Burlyanova (1960). When the content of silica and alumina in rivers, lakes, and seas reaches the necessary concentration, 110 authigenic aluminosilicates will crystallize. He cites essentially analcitic Devonian rocks that are not associated with volcanics. These beds are about 2.5 mm thick and composed of 90 percent closely packed analcime crystals. Evidence indicates these beds accumulated as an aluminosilicate gel that subsequently coagulated and crystallized to form a rock under quiet bottom conditions. Van Houten (1960; personal communication, 1963) believes the silica and alumina necessary for analcime in the Lockatong argillite of the Triassic Newark Group was available in true or colloidal solution. The zeolite may have been orig inally precipitated along with carbonate minerals, or it may have formed authigenically from a colloid or primary clay mineral. Milton and
Eugster (1959) report on fine-grained analcite in Eocene shales of
Wyoming. They say analcime forms by the action of saline waters on tuffaceous beds or by contact of sodic solutions and siliceous materials containing the necessary alumina.
Synthesis. —Synthesis of zeolites from feldspars and alkali carbonate solutions at 200°C and 15 atmospheres pressure was reported by Leonard (1927). Siren (1937) allowed Na(OH), Al(OH)3 , anc* silica gel to react in a steel bomb and found optimum temperature for produc tion of analcime to be 200°C. Spherulites of analcime were obtained by
Foster and Feicht (1946) from reaction of kaolinite and Na(OH) at 90°C and atmospheric pressure. The length of their run was 29 days. I l l
Furthermore, kaolinite was completely destroyed in a solution of NaCl and H2SO4. Coombs (1960) mixed kaolinite, calcite, and quartz in molar ratios of 1:1:7 and hy dr other mally produced wairakite, the cal cium analog of analcime. Kaolinite was altered to montmorillonite be low 273°C. Barrer and White (1952, in Deer, Howie, and Zussman,
1963) reported the hydrothermal synthesis of analcime from alumino silicate gels in the presence of excess Na(OH). Saha (1959), while studying the system nepheline-albite-water, found analcime as the only phase crystallizing from glasses of composition NagA^SigOiQ to
NaAlSigOg under hydrothermal conditions. He found that the cell edge decreases and HgO increases as SiOg increases.
Origin. —At Willcox Playa, negative evidence militates against considering devitrification of volcanic glass to form analcime. Further more, the occurrence of zeolite throughout the 140-foot core would re quire an almost continuous fallout of air-borne ash.
The formation of zeolite from solution requires a system sat urated with SiOg in the presence of quartz or amorphous silica. C hange in equilibrium conditions (evaporation) would precipitate colloidal silica which, under the proper conditions of alkalinity and sodium and alumi num concentrations, would form analcime. The solubility of silica at low temperatures has been studied by Krauskopf (1959) and Siever
(1962). Data germaine to this study are that silica exists in true
i 112 solution as monomeric H^SiO^, solubility is unaffected by pH in the range pH 2-9.5, and it is unaffected by other ions up to rather high con centrations. Amorphous silica is 10-15 times more soluble than quartz at 25°C, and thus one would expect amorphous silica to be dissolving and quartz precipitating in an aqueous system containing these two phases. This condition is not usually found in natural systems because quartz precipitates with great difficulty at low temperature (Siever,
1962) . Diatom frustules are a great source of amorphous silica in both the marine and lacustrine environments. Neither diatoms nor opaline silica were found in the playa sediments; therefore, we need consider only solutions of crystalline SiOg. Siever found the solubility of silica to be 10.8 ppm by extrapolation of the solubility curve to 25°C. He points out that the oceans are under saturated with silica, and that lakes, particularly sodium carbonate lakes, may have high concentrations of dissolved SiOg. Thus, Owens Lake, with 300 ppm Si02 and 213,700 ppm dissolved solids, would be ideal for crystallization of alumino silicate minerals such as zeolites. It is important that diatoms are known from these sediments (Benda, Erd, and Smith, 1960). Lake
Cochise was a sodium sulfate lake of much lower salinity. Analysis of a concentrated brine from a 6-foot observation well at the north end of the playa yielded 106, 000 ppm total dissolved solids (Brown, et al.,
1963) . Dissolved silica at 20°C was 9.7 ppm and alumina was nil. The presence of ostracods and the absence of evaporites in the core 113 sediments both point toward the existence of a brackish, rather than a highly saline, Pleistocene lake. As amorphous silica in the form of diatom frustules is not known from the sediments of Lake Cochise, a figure of 10 ppm dissolved SiOg for the waters of the lake during its active stages seems to be a fair interpretation of the geochemical data.
On this base, some calculations were made to determine if dissolved species could provide the necessary components to form analcime. In these calculations the ionic activity of dissolved SiOg is presumed to be equal to its concentration, and this is probably the case at such extreme dilution (G arrels, 1960). Now,
10 ppm = .001 gm/100 gmHgO - .01 gm /l, 000 gmHgO
= . 00017 form al, which means that in slightly less than 1 ,0 0 0 cc of water there is
0.00017 moles of SiOg. Analcime contains 55 percent SiOg and has a specific gravity of 2.25 gm/cc and forms about 10 percent of the sedi ment. With a porosity of 20 percent the bulk density of the sediment would be about equal to analcime. Therefore,
2.25 gm sediment/cc x 0 .1 gm analcime/gm sediment x
0.55 gm SiOg/gm analcime = 0.124 gm SiOg/cc sediment or, 0.002 moles SiOg are needed to form 10 percent analcime in each cubic centimeter of sediment. As the water contains 0.00017 moles of
SiOg per 1, 000 gm, a column of water 0.002/0.00017 x 1, 000 cm high would be required to provide the silica for each cubic centimeter of 114
sediment. This is 110 meters allowing for the greater density of saline water. Is this feasible? The present evaporation rate in the Willcox
area is 1.55 meters/year. Data are not available for lower evapora tion needed to sustain a lake, but at half the present rate it would re quire 140 years to yield SiOg for 1 centimeter of sediment. Martin
(1963) estimated the sedimentation rate in the middle of the lake to be
0.2-0 . 6 mm per year. He based his estimate on data from Searles
Lake, 0.1-0 . 6 mm per year (Flint and Gale, 1955); interpolation from
radiocarbon dates on the San Augustine Plain, 0.2-0.5 mm per year
(Clisby, unpublished); and from recent sedimentation in the deepest
portions of Salton Sea, 0. 5-0. 8 mm per year (Arnal, 1961). Using the
conservative sedimentation rate estimated for pluvial Lake Cochise, it
would require 50 years to accumulate 1 centimeter of sediment. This is
less than a third of the time calculated to precipitate the silica needed
for 10 percent authigenic analcime.
Admittedly these figures could be adjusted to a system favoring
the precipitation of analcime from dissolved species. It is considered
that the figures used are biased for such a system with the exception,
perhaps, of the analcime content in the sediment. This figure may be
high for some parts of the core by a factor of two. On the other hand,
this approach requires that waters replenishing the lake be saturated
with silica. Analyses of eight southwestern U. S. rivers (Clarke, 1924,
p. 87) give silica contents well below theoretical saturation values. 115
Streams draining into the playa probably contained even less silica, owing to the small area of the basin. Dissolved alumina has been ig nored in the calculations, and its presence may reduce the solubility of silica. This is probably due to the formation of an insoluble aluminum silicate on the surface of the solid SiOg phase (Cressman, 1962, p. 17).
Without discounting direct precipitation during dessication stages of the lake, the geochemical and petrographic data indicate that other sources of silica must be found.
Sand and silt fractions mounted in Caedax and gelatin (Nayudu,
1962) were studied with the petrographic microscope. Clean sands and silt were examined also with a binocular microscope against a black background. Quartz and K-feldspar showed no post-depositional solu tion effects, and ferromagnesian minerals showed "hacksaw" termina tions normal to these minerals from weathered zones. No secondary overgrowths were observed on any of these minerals. Plagioclase was a conspicuously minor constituent of the coarse fractions; however, the grains noted were angular and uncorroded. The lack of plagioclase may be significant but cannot be accounted for by diagenetic destruction.
Colloidal clay minerals offer the best source of silica and alumina for zeolitization. Kaolinite is unstable in an alkaline environ ment. All but a few of the nearly 200 pH values obtained from the core fell between pH 9.0 and pH 9.5 (fig. 16), and these compare favorably with direct measurements of hydrogen ion activity in other playa deposits 116
(Baas Becking, Kaplan, and Moore, 1960). Though not abundant, . kaolinite is being produced in the drainage basin under the present cli matic regime. The high pine pollen in the core from 6-69 feet is in dicative of higher precipitation and past conditions even more favorable for production of kaolinite around the playa.
Clay minerals soaked in sea water for 10 days yielded appre ciable amounts of SiOg, A^Og, and FegOg to solution (Carrol and
Starkey, 1960). Percent of the mineral lost by solution of these com plexes were in order: illite (.83), montmorillonite (.53), kaolinite
(. 51), mixed layer (. 22), and halloysite (. 20). Carrol and Starkey sug gest that the flow of silica from sediments to bottom water reported by
Goldberg and Arrenhius (1958) at all South Pacific stations may be pro vided by dissolution of clay minerals. A similar increase in SiOg was found by Emery (1960) in marine basins off southern California. He found that dissolved silica in interstitial water in the tops of cores was several times that in the overlying waters. Diagenetic loss of kaolinite is suggested by its complete absence in the core except for traces in the top ten feet. The formation of analcime from kaolinite involves a gain in silica relative to alumina represented by the formula,
Na(OH) + 2SiOg + AlgSigOg (OH)4---- ^(NaAlSigOg’ HgO) + HgO
sodium hydroxide + 2 silica + kaolinite—>analcime + water
On a molar-volume basis the amount of silica in kaolinite and analcime 117 is almost identical, but a volume-for-volume reaction would yield ex cess alumina. The extra silica needed could have been provided from alkaline waters or other clay minerals. The mineral relationships make it apparent that the sedimentary milieu was silica deficient.
Analcime has the lowest Si/Al ratio of high sodium zeolites (table 2), and on entropy considerations, Fyfe (in Coombs et al., 1959) considers it most unlikely that analcime would be more stable than mordenite at low temperature in a silica-rich environment. In the system NagO-
AlgOs-SiOg-HgO Hemley (1961) found that montmorillonite-analcime equilibrium is a function of the Na/H activity ratio. Montmorillonite forms at low ratios, whereas analcime forms at higher ratios. If the system were saturated with silica, then at times of reduced Na activity montmorillonite would form. It would be reasonable to expect also authigenic albite in such an environment. A plot of montmorillonite content versus analcime content showed no definable trend. One would anticipate a reciprocal relationship of montmorillonite formed at lower
Na activities or if analcime had formed from montmorillonite. Hay
(1963) found clinoptilolite, K-feldspar, and montmorillonite all forming from rhyolitic glass in the John Day Formation. Hay and Moiola (1963) suggest that analcime formed from montmorillonite in Searles Lake; however, they found only a tenuous reciprocal relationship between these minerals. The amounts of montmorillonite in the Willcox Playa core are not incongruous with that expected from the drainage basin. 118
The waters of pluvial Lake Cochise are not considered to have been excessively alkaline by the writer. According to Baas Becking,
Kaplan, and Moore (1960), lake waters with strong cations and anions, like sodium and sulfate, would have pH's near 7; those with a strong base (Na) and a weak acid (HCO3 -) would be alkaline as in trona lakes; and a strong acid (8 0 4 "^) with a weak base (Fe+^) would lead to a low pH. However, factors other than simple solution of soluble salts governed the waters of Lake Cochise. The high negative oxidation po tentials, black color, and fetid odor of the core are all indicative of stagnant bottom waters induced by meromixis. As dissolved matter became concentrated, heavier saline water settled to the bottom of the lake, and a stable density stratification was established. Metabolic processes and decay of organic matter depleted the dissolved oxygen at the sediment-water interface. Under such conditions anaerobic bac teria, particularly the sulfate reducing Desulfovibrio, would fluorish.
The sulfate in interstitial waters was organically reduced to HgS, some
of which reacted with iron to form the black pigmenting agent in the
sediments. The remainder rose to the sediment-water interface to
abet the already toxic bottom-water. The conditions outlined here have
been documented in other natural environments hundreds of times.
Emery (1960) found that in 400 measurements on 30 cores, 90 percent
of the samples with negative Eh’s emitted hydrogen sulfide. Further
more, he ascribes the rise in pH with depth in these sediments partly 119 to the reduction of sulfate by bacteria and to the generation of ammonia during oxidation of protenaceous and fatty organic matter.
In the near-surface interstitial waters of Lake Cochise, sulfate was reduced with a concomitant rise in pH and the Na/H activity ratio.
In this diagenetic environment conditions were favorable for dissolution
of kaolinite and crystallization of analcime. The decrease of sulfate ion
in interstitial water by bacterial reduction may be the major factor in
the absence of thenardite in the dried core samples.
Calcite
The factors affecting CaCOg deposition are temperature, partial
pressure of COg, pH, and other dissolved ionic species. In the system
CaCOg-HgO, calcite is precipitated below. 30°C, whereas aragonite
forms at higher temperatures. The latter polymorph may form at
lower temperatures in the presence of large divalent cations, such as
strontium and barium, that favor the more open structure of aragonite
(Zeller and Wray, 1956). The degree of saturation of CaCOg in natural
environments depends on these physical factors that are in turn affected
by biological activity. Atmospheric and metabolic COg in solution in
crease carbonate solubility by the formation of carbonic acid. Increas
ing temperature and photosynthesis reduce dissolved COg and decrease
carbonate solubility. In the pluvial lake, plants and high temperatures 120 tended to keep CaCOg saturation high, whereas ostracods served to re duce its concentration and degree of saturation.
The occurrence of authigenic calcite throughout the core in dicates that the waters of the lake were saturated during the time of sedimentation. Small crystal cluster, calcite accretions, and ostra- cod valves were found to a depth of 80 feet. Large calcite euhedra that form the bulk of the sand fraction occur from 80 feet to the bottom of the core. Conditions favoring increased calcite precipitation, probably an increase in mean annual temperature and evaporation rate, prevailed during the time represented by the lower 60 feet of core.
Evaporites
Thenardite (NagSO^), burkeite (NagCOg (8 0 4 )2 ), an<* halite a re found in white crusts on the playa surface. Thenardite and burkeite oc cur at Searles Lake as minor constituents of the edge and central
evaporite facies. Thenardite has been formed synthetically from pure
solutions of Na2S0 4 above 32.4°C. Burkeite occurs in the system
Na2S0 4 -Na2COg-HgO above 25°C (Palache, Berman, and Frondel,
1951). Neither of these minerals occurs in the dried core material.
This corroborates the writer's interpretation of the geochemical and
petrographic data of a lake with brackish waters and diagenetic loss of
sulfate in the bottom sediments. Neither thenardite nor burkeite would
be expected in a system high in Mg, Ca, or K, and their occurrence 121 supports the assumption that the brine analysis reflects the dissolved chemical species in pluvial Lake Cochise. GEOLOGIC SIGNIFICANCE
Climatic Episodes
Martin (1963) studied fossil pollen in the 140-foot core and in terpreted the vegetation history of the surrounding terrain in terms of
Pleistocene climatic changes. Taking high pine pollen as indicators of cool-wet climates and low pine counts or poor preservation as products of drier times, he recognizes several zones of paleoclimatic importance
(fig. 16).
(a) Zone 1, surface to 6.9 feet, contains no pollen. Probably due to oxidation under post-pluvial climatic regimes.
(b) Zones 2-4, 6. 9-69 feet, contains very high pine pollen.
Interpreted as a pluvial episode affected by Wisconsin glaciation in northern latitudes. The beginning of this climatic unit is placed at a depth of 75 feet.
(c) Zone 5, 69-140 feet, characterized by a lack of pollen from 75-95 feet that is interpreted as destruction by oxidation, perhaps during the Sangamon interglacial climatic episode. High grass and low pine percentages from 95-140 feet equated climatically with the Illinoisan
glacial stage.
122 123
I igure 16. --Clay mineralogy, pollen zones, and Eh-pH conditions in the core. 124
Martin allows other interpretations, namely that the entire core below 75 feet is Sangamon. It is not quite clear why he places the
Wisconsin-Sangamon climatic change at 75 feet, as this is a full 6 feet below the appearance of continuous high pine counts and pluvial condi tions.
A scrupulous comparison of the petrographic data with the vegetative history shows some interesting differences as well as areas of agreement. Large calcite euhedra appear at 80 feet and continue to the bottom of the core. They formed during a warm-dry or hot-dry episode during which the degree of saturation of CaCOg increased.
Partial dessication of the lake was indicated at depths from 113-117 feet and at 123 feet. This was manifest in abundant spongy, highly ox idized calcareous accretions occurring with unoxidized aggregates and clear authigenic calcite rhombs. Apparently marginal sediments were exposed to subaerial processes and then redeposited in a deeper part of the lake. Only one of these samples lacks pine in the pollen assemblage.
This is not surprising since the modern pollen rain on the dry lake bed contains 2-19 percent pine. There is no indication in the sediment record of dilution of Lake Cochise during a pluvial-glacial episode from
95-140 feet as proposed by Martin. Rather, progressively warmer con
ditions conducive to increased carbonate deposition are indicated by the
occurrence of larger and better developed euhedra through the interval
100-140 feet. 125
Eardley and Gvosdetsky (1960) report that a 650-foot core from
Great Salt Lake penetrated Pleistocene sediments equivalent to Aftonian interglacial through the last Wisconsin glaciation. Correlation with the classical Wisconsin glaciations and the Kansan glaciation are good, whereas two intervening pluvial (glacial?) cycles are incongruous with
Great Lakes chronology. They feel their chronology agrees better with the five cycles of stream entrenchment and alluviation that constitute the record of Pleistocene sea level changes in the Lower Mississippi
Valley. The top 30 feet of core is equated with the Wisconsin glaciation, preceded by an unnamed interglacial (Sangamon?) climate from 30-122
feet. There follows toward the bottom of the core two glacial-inter
glacial cycles to a depth of 428 feet, at which point deposits of definite
Kansan age are penetrated. The climatic interpretations are based on
petrographic and paleontologic criteria not unlike those of this study.
Interpolation from C-14 dates yielded an average sedimentation rate
of . 37 mm per year.
Flint and Gale (1958) recognized two lake stages in the upper
220 feet of core from Searles Lake, California. The mud deposits of
each lake were followed by precipitation of salt and near or complete
dessication. The latter of these pluvial cycles was 23, 000- to 10,000
B. P ., whereas the earlier pluvial stage was established by 46, 000 B. P.
and was waning by 32, 000 B. P. Flint and Gale suggest that the lakes 126
correlate with early and late Wisconsin glacial moraines from the Sierra
Nevada.
The Willcox Playa core is unlike either Searles or Great Salt
Lake in that clastic sediments dominate the section. There is no clean
contact between evaporites and muds as there is at Searles, nor is there
a preponderance of chemical and biochemical deposits as exists at Great
Salt Lake. It follows that the stratigraphic significance of near-homo
geneous deposits would be more difficult to decipher. On the evidence
from pollen analysis and carbonate deposition in Lake Cochise, a cycle
consisting of one pluvial and one interpluvial episode can be recognized.
Whether these are equivalent to a Wisconsin stade and interstade or to
• the Wisconsin glacial and Sangamon interglacial stages is not known.
Three carbon-14 analyses on core samples at shallow depths yielded
the following ages: A-221, 5-foot depth, greater than 20,000 years B.
P .; A-352, 6.25-7.0 feet, 23,000 + 500 years B. P .; and A-353, 7.7-
8.4 feet, 22, 000 + 500 years B. P. (Mr. Austin Long, personal com
munication, University of Arizona laboratory). Sample A-221 was on
organic material, whereas the other samples were on authigenic car
bonate crystals. Mr. Long states that because of the extremely small
size of the crystals and ground water movements at these depths, some
mixing of materials was possible. Taking the radiocarbon dates at face
value, the sediments of the 140-foot core represent a part of the
Sangamon-early Wisconsin climatic cycle of the Pleistocene Epoch. In 127 terms of age dates and sediment thicknesses, the Willcox Playa core agrees better with the bottom mud and mixed-layer zones from Searles
Lake than with any of the zones from Great Salt Lake.
It was anticipated that variations in clay-mineral content would reflect climatic changes. All attempts to find a systematic correlation between illite-montmorillonite ratios and other paleoclimatic criteria failed. There is a statistical increase in montmorillonite with depth, and it is tempting to associate this with progressively drier times in the drainage basin. The apparent increase is an artifact fabricated by the lack of well-organized montmorillonite and a high proportion of mixed-layers in the top 20 feet of core. Mixed-layers range from 15-30 percent of the clay minerals in these samples, whereas montmorillonite is usually less than five percent. It is believed that one or both of the following processes operated to produce the high proportion of mixed- layers.
(1) Pollen counts show as much as 99 percent pine in this zone reflecting cool-moist conditions promoting the formation of degraded or
"stripped" illites. Upon entering the lake the degraded layers adsorb sodium, calcium, and magnesium with concomitant expansion and de velopment of interstratification. The organization of montmorillonite was diminished by the strong dispersing action of sodic waters. This latter phenomenon has been used to explain the poorly organized 128 montmorillonites in the bottom sediments of Great Salt Lake (Grim,
Kulbicki, and Carrozi, 1960).
(2) Oxidation potentials indicate that at least the top eleven feet of sediment has been esqposed to the atmosphere and percolating waters. Very fine-grained montmorillonite, less than 0.1 micron when dispersed (pi. 1) was selectively transported by downward moving waters to the water table. At the same time illite was leached with the formation of mixed-layers. The very low content of montmor illonite in all the playa surface samples lends credence to such a process.
The writer concludes that the clay minerals of the Willcox
Playa core are of little paleoclimatic significance. Eardley and
Gvosdetsky (1960) reached a similar conclusion about the Great Salt
Lake core. They generalize that it is not safe to interpret fresh and saline environments from the clay minerals present. In the present state of the science the writer would agree. It is necessary to study the whole mineral assemblage in order to make competent judgments about past events.
Diagenesis
Definition
Most geologists agree, in a general way, that diagenesis is the physical and chemical changes wrought on sediments in the depositional 129 environment. Disagreement lies in delineating the change-producing conditions and processes of the diagenetic milieu from those of meta morphism, on the one hand, and weathering, on the other. Three of four definitions in the “Glossary of Geology" (Howell, 1957) allow that diagenetic changes are those that take place in sediments before lithifi- cation or consolidation. It is generally understood that these changes take place at temperatures and pressures near those at the earth’s sur face. Some authors include post-lithification changes provided these changes are not katamorphic in character (Whitehouse and McCarter,
1958). Thus biological activity, Eh, and pH are most important in syngenetic and early diagenetic reactions, whereas temperature and pressure play subordinate roles. Temperature and pressure become increasingly important with depth of burial and may drive epigenetic reactions, and, of course, they are the critical factors in metamor phism. As Packham and Crook noted (1960), Eh and pH cannot change the lines on a phase diagram, but they can prevent a phase from form ing all together. Diagenetic reactions are a response of the assemblage toward equilibrium, and in this respect they are analogous to metamor phism. It is " .... after all, but the introduction to metamorphism"
(Sujkowski, 1958). 130
Diagenesis of clay minerals
In recent years clay mineralogists have found it necessary to pledge allegiance to one of two schools of thought regarding clay mineral diagenesis. Prior to 1958 diagenesis in the marine environment was a generally accepted phenomenon. Grim, Deitz, and Bradley (1949),
Grim (1953), Powers (1954), and Grim and Johns (1954) had demon strated that illite and chlorite form from less well-crystallized material and that some loss of kaolinite was indicated. Weaver (1958a, b, c,
1959) challenged this concept and cautioned that changes controlled by the detrital nature of the clay mineral are not diagenetic. Cation ex change or cation fixation is not diagenesis because the basic structure of the mineral has not been altered, " .... but because different names are attached to the clay before and after exchange, the process is called diagenesis" (Weaver, 1958a). He documents this paper with numerous examples of reported "diagenetic" changes that are simple cation-ex change processes, such as the fixation of K* by montmorillonite to form mixed-layers and the uptake of Mg+^ by degraded chlorite to form better crystallized material. Weaver concludes that: (1) clay minerals do not originate in their depositional environment, and (2) generally
they are not strongly modified by depositional environment. He extends
this generalization to bentonitic montmorillonite and glauconite empha
sizing that the mineral crystaUized depends on the parent material 131 rather than the depositional environment. A number of papers appeared subsequently that attribute the distribution of clay minerals in the west ern Atlantic Ocean and the Gulf of Mexico solely to physical phenomena, i. e., source area, current sorting, and differential flocculation (Pryor and Glass, 1961; Milne and Earley, 1958; Griffin, 1962). If cation ex change processes are excluded from diagenesis, then the area of dis agreement amongst mineralogists is greatly reduced. It is Grim's
(1958) opinion that there would be less difference of opinion if there was more agreement on the analytical methods upon which diagenetic conclusions are based.
In the marine environment cations will be adsorbed (exchanged) in order of their free bonding energy, Mg+> Ca+! Na+, to sites with the highest bond energy (Carrol and Starkey, 1960). Therefore it would be predicted that degraded chlorites take up magnesium. According to
Weaver (1959) only minerals with tetrahedral charge will fix potassium, e. g., stripped illite or beidellite, and therefore no structural change is involved. However, in order to accept the thesis that clay minerals are little effected in the depositional environment it is necessary to ignore many field observations as well as laboratory work on synthesis and phase equilibria. Nelson (1960b) cannot explain the chlorite increase along the Rappahannock Estuary salinity gradient by cation exchange processes. Diagenetic illite and glauconite have been demonstrated in the Morrison Formation on the Colorado Plateau (Keller, 1960). The 132
ease of formation of chlorite from montmorillonite has been demon
strated in the laboratory (Whitehouse and McCarter, 1958; Slaughter
and Milne, 1960). The whole field of mineral equilibria is predicated
on the fact that minerals are more "at home" in some environments
than in others, and clay minerals are no exception.
Willcox Playa sediments
It is the writer's opinion that loss of kaolinite to form analcime
. has occurred. The evidence and inferences for such a transition are
presented below.
(1) There is a minor amount of kaolinite forming in the source
area today, and conditions were more favorable for its development
during pluvial periods, yet this mineral is essentially absent in playa
sedim ents.
(2) In the absence of diatom frustules and amorphous silica,
it is considered that the low solubility of quartz would be insufficient to
provide the necessary constituents. Saturation would be reached only
at times of dessication because dissolved SiOg in streams of arid re
gions is characteristically low (Clarke, 1924).
(3) There is no evidence of dissolution of other silicate min
erals, although plagioclase is anomalously low in the sediments.
(4) The high pH and relatively low salt concentration of inter
stitial water of the playa sediments are conditions particularly unsuitable 133 for kaolinite. Other clay minerals may have contributed silica and
alumina to interstitial solutions, but there is not evidence that they were
destroyed_hi situ.
The instability of kaolinite in saline environments is recognized,
and its loss in marine deposits has been indicated. Millot (1942, in
Grim, 1958) indicates the loss of kaolinite in marine waters would be
particularly high if carbonate were also accumulating. At Boron,
California, kaolinite is found in red siltstones but is not associated
with green beds and borate deposits. Gates (1959) suggests that high
pH destroyed the kaolinite in the barren beds. Grim (1958) states
M.... in highly saline lakes the two-layered minerals would disappear, ”
and further on in the same work, " ... .this matter of possible loss of
kaolinite in alkaline environments deserves intensive study, and seems
to be amenable to laboratory investigations."
Against the dissolution of kaolinite in saline waters is the fact
that this reaction has never been unequivocably demonstrated. Droste
(1959, 1961a, b, c), aware of the possible loss of kaolin in California
playa sediments, concluded that in no case did a clay mineral coming
into a playa disappear in the depositional environment. Furthermore,
he found the clays of the playa surface to be just what was being trans
ported to the lake. He found kaolinite in 13 out of 45 playas, including
Searles Lake. In the long core from Searles Lake analcime occurred
in 21 samples, six of which contained a 7A mineral. In their study, 134
Hay and Moiola (1963) assigned all the 7A reflections to kaolinite whether chlorite was present or not. They based their identification on Droste's
(1961b) recognition of kaolinite from the same core; however, Droste admitted kaolinite was possible but not positively identified. The ex istence of kaolinite in the saline section of the Searles core is by no means firmly established. The chlorite-kaolinite ambiguity is a prob lem faced by all clay-mineral workers. No good techniques have been developed as yet to distinguish these minerals in small amounts. In the
Willcox Playa core its absence was easy to establish because 7A reflec tions disappeared upon mild heat treatment of glycolation. In many in stances 7A reflections were totally missing.
The absence of kaolinite cannot be readily explained by sedi mentary dispersal mechanisms or source area. The total mineralogy of the core, insofar as the writer has studied this material, is best ex plained by diagenetic destruction of kaolinite and its reorganization to form analcime.
Par agenesis
Lack of textural evidence makes it difficult to determine para-
genetic relationships. The geologically recent nature of this deposit
limits observed changes to syngenetic (halmyrolytic) or early diagenetic
processes. Much of the calcite in the sediments occurs as isolated
euhedra much larger than pore spaces in the compacted sediment and 135 therefore represents a direct precipitate from bottom waters. The same is apparently true for aggregates, pseudo-oolites, and faecal(?) pellets. The tiny birefringent specks of calcite seen in mounted sec tions may have formed from interstitial waters, but this alternative is not necessary to explain the origin of this mineral.
The data cited from marine sediments and laboratory experi ments and the chemical state of the core indicate that conditions for analcime crystallization were optimum in interstitial waters. Condi tions of high alkalinity, Na/H ratios, and silica concentrations would have been obtained a short distance below the sediment-water interface.
It is doubtful if Eh exerts any control on analcime formation.
Mineral Facies
A .metamorphic zeolite facies was described by Fyfe, Turner, and Verhoogen (1958) to fill the gap between low-grade green schist facies and diagenesis. Sedimentary rocks may be altered on a regional
scale, without the aid of deformation, to mineral assemblages resembling
those of diagenetic associations, and these should be considered meta
morphic. Justification for such a facies was based on extensively zeo-
litized polymictic sediments from New Zealand that contained quartz-
laumontite-albite and quartz-albite-pumpellyite-adularia assemblages.
From consideration of zeolite paragenesis, some of which are by no
means metamorphic, Coombs (in Coombs et al., 1959) extended 136
Turner's metamorphic facies to a zeolite mineral facies in the sense of
Eskola's mineral facies concept. Thus the zeolite facies was general ized to include assemblages produced under physical conditions in which quartz-analcime, quartz-heulandite, and quartz-laumontite are com monly formed, whether products of diagenesis, hydrothermal activity, or conventional metamorphism. They caution that facies designation, to be of value, should be restricted to regionally developed assemblages.
Such occurrences as interstitial aggregates and cements may be called zeolite facies in the informal sense of Eskola.
Russian geologists have applied the concept of geochemical facies to sedimentary associations with much success. Teodorovich
(1949, reviewed by Chilingar, 1955) subdivided subaqueous depositional environments on the basis of pH-Eh conditions. The sediments of
Willcox Play a belong to Teodorovich's soda facies, delineated by oxida tion potentials conducive to formation of HgS or FeS, and pH greater than nine. Rengarten*s (1950, in Coombs et al., 1959) geochemical zeolite facies is characterized by alkaline conditions, colloidal solutions of SiOg and AI2O3, and high oxidation potentials. According to
Rengarten, oxidizing conditions separate the zeolite facies from the glauconite facies. The results of this work and many of the other oc currences cited show that zeolites may form under either oxidizing or reducing conditions. 137
Krumbein and Garrels (1952) qualitatively classified marine sedimentary environments and sediments in terms of Eh and pH. They constructed a diagram in which mineral assemblages were "fenced" into areas delineated by oxidation-reduction potentials and pH. The sediments of Lake Cochise fall along the sulfide-sulfate fence with alkalinities well above those necessary for carbonate deposition. This is to be expected as the bulk of chemical sediments form to the right of this fence. The range of Eh-pH values is delineated in figure 17. The results of laboratory synthesis and phase equilibrium studies cited pre viously corroborate the natural conditions as measured in these sedi ments. It is apparent that crystallization of analcime at low tempera tures is a function of Na/H ratios with little control by Eh.
Packham and Crook (1960) suggest the possibility of erecting a sequence of diagenetic facies analogous to metamorphic facies from the surface down to metamorphic zones. They define such facies as " .... including all rocks or sedimentary detritus which have, by a process of diagenesis, developed mineral assemblages that are the result of adjustment to a particular diagenetic environment." Parameters of a diagenetic facies would be temperature and pressure, with Eh-pH pa rameters as subfacies called "parfacies. " Their facies do not differ
much from Coombs zeolite facies except that equilibrium conditions
are not required and it is applied to sedimentary rocks. They consider
that the changes manifest in a rock by temperature, pressure, Eh, and 138
8.0 9.0 *• *
8.0 ligure 17. --Diagenetic geochemical environment of the Willcox Playa to a depth of 140 feet. (After Krumbein and Carrels, 1952.) 139 pH are diagenetic changes until the fabric of the original rock is exten sively changed. They recognize a heulandite-analcite diagenetic facies equivalent to an Eskola-type mineral facies and equal to the conditions existing in the upper 15, 000 feet of geosynclinal sediments described by Coombs (1960) from New Zealand.
The geochemical facies of Teodorovich, or the fence diagram
of Krumbein and Garrels, is most applicable in describing the environ
ment at Willcox Playa. It seems imprudent at the present time to char
acterize such accumulations by a "facies” designation in view of
inadequate knowledge concerning what minerals should exist at equilib
rium with saline waters. SUMMARY OF CONCLUSIONS
1. The clay minerals of the core are, in decreasing order of abundance, illite, montmorillonite, vermiculite, and chlorite. They reflect the clay minerals forming in the source area and are of detrital origin. Kaolinite occurs only in trace amounts in the top few feet of core material.
2. No Pleistocene paleoclimatic significance is attached to variations in the ratio illite: montmor illonite or illite: montmor illonite +
vermiculite. It is concluded that the clay minerals in the core cannot be used to deduce changes in the salinity of pluvial lake Cochise.
3. Authigenic analcime constitutes 5-10 percent of the clay
fraction of the 140-foot core. The absence of volcanic ash and low
concentration of silica in brines close to the playa surface preclude
reorganization of felsic glass or crystallization of precipitated alumino
silicate gels to form analcime. Consideration of the high pH values and
the whole mineral assemblage of the core favors the diagenetic dis
solution of clay minerals, probably kaolinite, to provide the necessary
silica and alumina for analcime crystallization.
4. The low percentage of authigenic calcite, the occurrence
of ostracod valves, and the absence of evaporite minerals in the core
140 141 records a history of brackish-water conditions for pluvial Lake Cochise.
5. The heavy minerals in the core sediments are predominant ly those expected from rocks to the north of the present dry-lake bed.
6. Real differences exist between the clay minerals of the playa surface and those in stream beds and soils in the source area.
The low proportion of expansible minerals at the playa surface is best explained by selective removal by downward percolating waters, ab sence of sediment contributed from areas high in montmorillonite both east and west of the playa, and perhaps deflation of finer colloidal ma terial at the dry-lake surface.
7. The pollen profile, radiocarbon ages, and history of car bonate deposition in the cored interval may be interpreted as represent ing one dry-warm and moist-cool Pleistocene cycle, perhaps equivalent to late Sangamon-early Wisconsin climatic episodes. APPENDIX A CONTRIBUTION OF VARIOUS PARTS OF MICA LAYER TO THE STRUCTURE FACTOR OF BASAL REFLECTIONS
Section of Composition F(10A) F(5A) F(3.3A) structure
Anion (Si, Al)80 2o(OH)4 17 120 74 skeleton
Interlayer 2. OK 37 34 M 1 .5K 0 .5H20 32 30 26 is is is 1. OK 1 .0H2O T l 25 0 .5K 1 .5H20 23 21 2.0H2O 18 16
Octahedral 4A1 45 42 39 3A1 Fe 58 54 50 2A1 2Fe 71 66 60 1A1 3Fe 85 78 71 4Fe 98 90 91
6Mg 68 62 56 5Mg Fe 81 75 67 4Mg 2Fe 94 87 78 3Mg 3Fe 107 99 89 2Mg 4Fe 120 111 100 IMg 5Fe 133 123 111 6Fe 146 136 122
For layer structures the F*s of basal reflections are found by
F(00JZ ) = 21 fa cos2-n.fi z
and for dioctahedral micas may be evaluated by the partial sums
142 143
4f (Al, Mg, Fe) + 2f(K, H20) (-1)^ + 8f(Si, Al)cos2 iri z (SiAl) +
12f0(cos2 Tri z0 l + COS2tcJLz 0^)
As the F contribution from the anion network is essentially constant for a given reflection, one need only evaluate changes in F for variation in octahedral cations. layer structure factors from Brown (1955). APPENDIX B HEAVY MINERALS OF THE CORE
£ Other Minerals— o Comments
I Tourmaline Chlorite E-C-Z Biotite Garnet CO Rutile Apatite
203 X Very poor suite. 208 r r r r VC s VC Basaltic hnb. pres ent. 212 r VC s VC Sphene(r), basaltic hnb. 216 X X X Very poor suite. 17 c VC c 19 X XX X Very poor suite. 25 c c VC 34 No heavy minerals. 43 X X Very poor suite. 46 No heavy minerals. 51 c c VC Fluorite. 55 VC Fluorite. 59 No heavy minerals. 69 X Fluorite(x), olivine (x). 78 X Very poor suite. 86 c s c c Very fine grained. 93 c c c Fluorite(r), bas. hnb. 106 c c s Fluorite(c), bas. hnb. 114 s s c c 118 s c c s c Basaltic hnb. 125 r s s VC 135 r r c c VC Bas. hnb., sphene (r). 139 r r VC s Poor suite.
144 145
Other Minerals— Comments Zircon Pyroboles Tourmaline Sample No. Depth (feet) E-C-Z Garnet Chlorite Rutile Apatite Biotite
142 104 VC r r C s s c Bas. hnb., sphene (r). 147 108 VC s r r r s c VC 150 109 No heavy minerals 155 113 VC r s c c VC Green tour., bas. hnb. 161 117 VC r r s c c c Bas. hnb. 166 120 c r s c c Poor suite. 172 123 VC r r r c VC 177 126 VC r r r c VC 187 132 a r r s s c 195 137 VC r s c c Brown tour., bas. hnb.
a = 50-75% vc = 25-50% c = 10-25% S r 10-2% r = few grains x = too few grains present for statistical estimate LITERATURE CITED
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