Chapter 3 Diagenesis of Sandstones and Compaction

Chapter 3

DIAGENESIS OF SANDSTONES AND COMPACTION KARL H. WOLF and G.V. CHILINGARIAN

GENERAL FACTORS CONTROLLING COMPACTION OF SANDSTONES

Any discussion on the origin of sedimentary rocks should include references to compaction, lithification, and other diagenetic processes as shown in the diagrams of Figs. 3-1 and 3-2.These diagrams demonstrate that compaction is one process that determines the final property of a rock, and close scrutiny of Fig. 3-1,for example, will show that any primary sedimentological factor that controls the original characteristics of the sedimentary deposit

J I

I +

I

L_________________-__----------------J

Fig. 3-1. Theoretical system diagram showing principal processes that affect shallow-water marine sedimentation. Dashed outline separates exogenous processes that supply inputs to the system but do ?ot receive feedback from it. Endogenous processes are inside dashed outline. (After Harbaugh and Merriam, 1968; in Harbaugh and Bonham-Carter, 1970, fig. 7-1, p. 265; courtesy Wiley-Interscience, New York.)

70

K.H. WOLF AND G.V. CHILINGARIAN

Uplift

I

Fracturing

I

!

Fig. 3-2.Diagram of fluvial system. Uplift and climatic factors of temperature, precipitation, and wind are outside the system, feeding into the system, but not receiving feedback from the system. Lithification and compaction are also outside the system, but receive output from the system instead of feeding into the system. (After Harbaugh and Bonhamcarter, 1970, fig. 1-14, p. 22; courtesy Wiley-Interscience, New York.)

will also determine the compaction history of the accumulation (see Introduction chapter). A possible sequence of interdependent, large-scale geologic parameters, which influence compaction directly or indirectly, has been given in Fig. 3-3. Purely theoretical considerations or, as in the sequence in Fig. 3-3, mere “common sense”, may be sufficient to think of other similar mega-relationships. In the future, however, more subtle and quantitative interdependencies must be determined t o make progress in solving complex problems in sedimentology, as has been attempted by Wolf (1973a,b),

TECTONIC-~AO~~~~~~~~,~~-OR EXTERNAL INFLUENCES

A C C u M U L A T I O ~ ~ ~ ~ COMPACTION ~ ~ ~ ~ ~ - , EROSION C-DIAGENESIS

ADJUSTMENT

Fig. 3-3.Flow diagram showing the relationships between various large-scale geologic variables during the development bf a basin and sedimentation that influence compaction diagenesis. These, in turn, have a reciprocal control on the large-scale variables.

DIAGENESIS O F SANDSTONES AND COMPACTION

71

among others. To do this, it seems best to start not with the regional, but with the microscopic and mesoscopic processes and factors involved in compaction. These concepts should then be applied on a regional scale. Review of the literature shows that a large amount of data on compaction is already at hand, especially on the more local scale. The information on the vertical and horizontal variations of the effects of compactional diagenesis, on the other hand, is somewhat meagre, but several examples are available from the literature. In compaction studies of sandstones many factors are to be considered (Table 3-I), many of which are related to the presence or absence of components additional to the sand-sized fraction. For example, the compaction history of a well-sorted sandstone is different from that containing clay matrix. Also, a stratigraphic section composed of sandstone alone may have a much different history of compaction as compared with a section composed of sandstones interbedded with siltstones and/or shales. In an investiTABLE 3-1 Factors and processes controlling compaction of sandstones Group I (Inherited factors1 : individual properties of grains and fluids)

Group II (Inherited factors’ : mass properties)

grain size grain shape grain orientation grain surface features grain rounding grain sphericity grain electrostatic properties grain composition impurities on grain’s surface fluid composition

grainlmatrixlcement ratios shalelsiltstonelsandstone ratios in stratigraphic section absorbability adsorption capacity degree of cohesion size sorting compositional sorting thixotropic properties porosity and permeability grain stacking patterns packing packing heterogeneity grains’ frictional properties grains’ total surface area shear strength surface area configuration of deDosit topography (e.g., slope) of unit sedimentary structures

Group III (Dynamic factors) rate of fluid movement replenishing of fluid removal of fluid fluid pressure overburden pressure subsurface temperature rate of sedimentation (= rate of loading) earthquakes tectonism time Group IV (Inhibitory factors’ : reducing compaction) diagenetic changes neomorphism cementation recrystallization epigenesis (catagenesis) metamorphism

This term was adopted from Coogan and Manus, see Vol. I, Ch. 3.

72

K.H. WOLF AND G.V. CHILINGARIAN

TABLE 3-11 Factors and processes significant in controlling compaction of pyroclastics' Factors

Remarks

(1)Heat retained in the deposit

may cause welding of the particles, thus increasing the strength of the deposit controls the rate of increse in overburden

( 2 ) Rate of settling and accumulation of the pyroclastic debris (3) Distance from source and thus differential sorting of pyroclastic debris (4) Degree of reworking in the depositional environment (5) Amount of rainfall during volcanism (water precipitation is often reported during volcanicity) (6)Amount of ground water

controls thickness of deposit and, therefore, overburden pressure controls textures, fabrics, and structures of deposit controls initial compaction

influences diagenesis, e.g., controls cementation by zeolitization, which increases strength of deposit ( 7 ) Composition (mineralogic and bulk chemi- controls degree of reactibility (or chemi. cal) : cal stability) during diagenesis, as well as (a) acidic versus intermediate versus physical strength prior to and subsequent basic volcanic particles to compaction (b) ratio of phenocrysts/fine debris (c) bubbles versus groundmass in shards determine degree of resistance to mecha(8) Shape, internal textures, and fabrics of vitric shards nical compaction

Most or all of the factors listed in Table 3-1 are applicable here also, but have not been duplicated. Only those of specific importance in pyroclastics have been listed here.

gation of sandstone compaction, therefore, no clear separation can be made between the various lithologies, as illustrated by the many interdependent parameters in Table 3-1. It is also obvious that certain factors of importance in the compaction of pyroclastics either need not be considered in epiclastic (= terrigenous) sediments or are of relatively minor significance (Table 3-11). It may be advantageous to briefly explain compaction of sediments by using the model offered by Terzaghi and Peck (1968; see also Dickey, 1972) and illustrated in Fig. 3-4. A cylinder full of water, containing several perforated plates which are separated by springs, is used in the experiment. The small holes in the plates represent the low permeability of clays, whereas the springs imitate the strength of the clay-mineral accumulations that can sustain the weight of the burial or overburden pressure. At various levels between the plates, the cylinder is connected to manometers measuring the pressure of water. At the start of the experiment, i.e., before loading,

DIAGENESIS OF SANDSTONES AND COMPACTION

73

Fig. 3-4. Terzaghi's model of the compaction process. When a load is applied to the cylinder, the water first carries the full load. As water escapes through perforated plates, the springs sustain the load and water pressure drops t o normal hydrostatic. At time t = 0, pressure (p) is excess hydrostatic; pressure head in manometers ( h l ) in feet is equal to plyw, where p is in lb/sq. ft and 7,,, is the specific weight in lb/ft3. At time t l , pressure drops in the upper part of the model (first three manometers). At time t 2 , pressure drops in all four manometers. Finally, at time t = O 0 , the pressure is normal hydrostatic (level of fluid in all four manometers is at the top of model. (After Dickey, 1972, fig. 3, p. 6; courtesy Int. Geol. Congr., Montreal.)

-

the water in the manometers is at the same level as that in the cylinder (coincides with the horizontal line marked t = in Fig. 3-4).If pressure is suddenly applied to the cylinder (representing overburden pressure), the springs will compress and water will escape from between the plates. As the holes are small, the movement of the fluid will be slow so that compression of the springs is initially prevented and the load is carried by the water itself, leading to excessive hydrostatic pressure. This stage is presented by a horizontal line (t = 0), indicating that the water is at the same level in all the manometers. Subsequently, the fluid will pass out of the upper part of the cylinder and the pressure will decrease successively as time passes from the upper to the lower part, as indicated by the curves t l and t 2 to t,, until the hydrostatic pressure becomes very small and the full weight of the load is carried by the springs. The springs, in this case, represent the framework of the rock which would carry the overburden pressure. In sedimentary basins, the loss of large volumes of water takes a very long time, and excess hydrostatic pressures are common in younger formations, as described in numerous publications. One example, offered by Kidwell and Hunt (1958),is presented in Fig. 3-5: in the Recent sediments of the Orinoco delta at Pedernales, Venezuela, excess pressures increased to 8 psi and above at depths of 120 f t below the surface. The pressure decreased upward to zero at a permeable sandy bed, which allowed quick passage of intrastratal fluids,

74

K.H. WOLF AND G.V. CHILINGARIAN PCH-9

PCH-I1 20 L

60

% I l-

100

a

g

140

Fig. 3-5. Excess hydrostatic pressures (numbers in psi) in the recent sediments of the Orinoco River, Pedernales, Venezuela, As the muds compact, pure water is bleeding upward to a permeable bed at depth of about 30 f t , and also downward to the pre-Paria unconformity, which is also permeable laterally. (From Kidwell and Hunt, 1958; in Dickey, 1972, fig. 4,p. 6 ; courtesy Int. Geol. Congr., Montreal.)

and downward to about 4 psi at an unconformity. The latter permitted bleeding off of the water. DIAGENESIS IN GENERAL*

The widely-accepted division of sedimentary processes into syngenesis, diagenesis, and epigenesis is inadequate for detailed research. The stages at which most of the individual processes and products occur cannot be distinctly demarcated. Two or more processes may be active simultaneously. They may overlap or the termination of one may mark the commencement of another process; and still other alterations may occur independently in both space and time. Many of the interpretations depend on the scale - thin section, hand specimen or outcrop -- at which observations are made. Hence, “pigeonholing” of processes without contradictions is difficult, sometimes even impossible. Difficulties in genetic interpretations occur in particular in mono-mineralogic rocks, such as fine-grained quartzites. For these and other reasons, it is not surprising that no agreement has been reached on the definition and extent of diagenesis (see review by Teodorovich, 1961). Diagenesis has been restricted to those processes that cause lithification. *In dealing with the ge’neral aspects of diagenesis of sandstones, the section presented here has been adopted with minor changes from Chilingar et al. (1967).

DIAGENESIS OF SANDSTONES AND COMPACTION

75

Such a limited application, however, is arbitrary, artificial and impractical.

Not only are there several distinctly different lithification processes which are frequently difficult to recognize and separate, but they are so gradational as to defy precise definition and can occur at any stage during the early history of sediments. It is virtually impossible, therefore, to exclude from

diagenesis other early alteration. It is preferable to apply diagenesis in a wider sense to processes that affect a sediment after deposition and up to, but not beyond, lithification andfor filling of voids. Although these latter processes can, and usually do, take place at different times within a sedimentary formation, especially if composed of different facies, the final stage of lithification and/or filling of voids appears to be the most convenient time at which diagenesis can be terminated. * Hence, the following rather all-inclusive definition, in general agreement with the concepts of Ginsburg (1957) and Krumbein (1942),has been adopted here: diagenesis includes all physicochemical, biochemical, and physical processes modifying sediments between deposition and lithification at low temperatures and pressures characteristic of surface and near-surface environments. Postlithification processes grade into epigenesis, and epigenesis passes into metamorphism.** ’Epigenesis near the depositional environment is called juxta-epigenesis (“juxta-” meaning near), and epigenesis remote from the surface is named apo-epigenesis*** (“apo-” meaning far, remote). Most coarser-grained sediments have some small voids which have been partly or wholly filled by one or more generations of cement. Hence, it is possible in some cases to divide diagenesis into pre-, syn-, and postcementation stages. In other paragenetic investigations, however, the diageneticepigenetic boundary may have to be based on some other criterion to be determined by the individual investigator concerned. No definite rule is possible. As long as the boundary is precisely defined by certain fabric or structural relations, little confusion should occur. A diagenetic-epigenetic *This approach has been found to be of particular use in coarse-grained or open-textured sandstones, conglomerates and limestones, but may be more difficult to apply to finegrained rocks. Nevertheless, the paragenetic model used here is convenient, because after lithification (= infilling of voids by cement) of a rock, the intrastratal fluids related to surface conditions cannot penetrate readily the rock framework. In cases where sediments maintain their porosity and permeability for a long period of time, even after being far removed from the original depocenter, the intrastratal fluids occupying the cavities can also be looked upon as either of syngenetic, diagenetic or epigenetic origin. **Strictly speaking, epigenesis as defined here, passes into metamorphism only if an increase of pressure and/or temperature occurs. ***Weathering not related to the original depositional environment of the sediments is not included in diagenesis and epigenesis as defined here. The gradational frontiers between deposition (= syngedesis), metamorphism, and weathering have been illustrated by Dunoyer de Segonzac (1968) in Fig. 3-6.

K.H. WOLF AND G.V. CHILINGARIAN

76

boundary established on the basis of a few thin sections of a local outcrop, however, may have t o be revised and shifted up or down the paragenetic scale as soon as the petrologic and petrographic information of the whole formation is available. On the other hand, it may be found that the termination of diagenesis in one area may be completely unrelated to that of other localities. Not all diagenetic stages are present in sediments. For example, precementation dolomitization, with the formation of dense and relatively impermeable dolomite, may completely alter the calcareous components of a limy sandstone, resulting in an absence of the syn- and postcementation stages. Postcementation dolomitization may obscure the various cementation stages. The raw material of diagenesis, as Krumbein (1942) called it, consists of organic and inorganie sediment of allochthonous and/or autochthonous origin, interstitial fluids, and other components subsequently formed or introduced into the system. In general, it is possible t o subdivide the components that interact during the diagenetic processes into the following (Wolf, 1963): (1) diagenetic-endogenic; (2) diagenetic-exogenic: (a) supergenic-exogenic; (b) hypogenic-exogenic. This is merely an expansion of Amstutz’s (1959) division: syngeneticsupergenic, syngenetic-hypogenic, epigenetic-supergenic, and epigenetic-. hypogenic. In most cases, diagenesis derives its raw material from both endogenic (within the sediments) and exogenic-supergenic (outside source-from above) sources. One or the other may prevail. Under unusual conditions, however, a volcanic, i.e., exogenic-hypogenic, source may supply components for diagenesis without a marked increase in temperature. This would be particularly true for siliceous material introduced into a geosyncline (or other depocenter); during diagenesis of the sediments, quartzose arenites can WEATHERING

DEPOSITION

METAMO~~~ISM

Fig. 3-6. Diagenetic frontiers. (After Dunoyer de Segonzac, 1968, fig. 1.)

DIAGENESIS OF SANDSTONES AND COMPACTION

71

be converted t o orthoquartzites, carbonates can become siliceous, and fossils can be replaced prior to dolomitization. Diagenesis may express itself in a number of different ways. Krumbein (1942) mentioned that a total of about thirty separate diagenetic processes have been described in the literature. They may result in mineralogical changes, addition and removal of material, and textural and structural modifications and alterations ranging from slight to extensive or complete. Several generations of diagenesis may each leave evidence, or each successive one may obliterate or destroy the products of earlier processes. In many cases, however, diagenesis may appear to be absent if only visually-obtained information is considered. * The division of sedimentary processes into several subdivisions is suitable to our present purpose and state of knowledge of sandstone diagenesis, including compaction. It should be pointed out, however, that the pre-, syn-, and postcementation sub-groups will become arbitrary and artificial with an increase in our understanding of diagenesis. To ascribe to certain products merely a genetic term such as “precementation-diagenetic”, without relating it to the sediment’s history as a whole, invites criticism. It is more accurate to relate all processes and products t o a paragenetic sequence. In other words, a paragenetic scheme furnishes less ambiguous information in cases where it seems impossible to define exact syngenetic-diagenetic-epigenetic boundaries. The absolute time of formation may be impossible to determine, but the textural and structural relationships permit the interpretation of relative time of formation. The widely used terms “primary” and “secondary” have very little meaning in diagenetic investigations unless precisely defined, although they may be quite useful in a very general colloquial sense. Paragenetic interpretations are relatively easy and noncontroversial if the investigation is made on the scale of one thin section or hand specimen. Syn-, dia-, and epigenetic processes, however, are not only gradational, and overlap in time and space on a microscopic scale, but especially do so on a regional scale. Regional diagenetic studies may be rather tedious and resemble structural analyses, for example, in that micro-, meso-, and macroscopically examined features are assembled step by step. The so-called predepositional (or presyngenetic) processes and products can be deduced from rock fragments, which were derived from older sedi*It should be noted that diagenesis is not part of the petrographic (= descriptive) stage, but belongs to the subsequent stage of petrology and petrogenesis (= interpretive), although diagenetic data is often made part of petrographic descriptions. Reliable diagenetic reconstructions cannot be made, therefore, on the basis of the study of a few local thin sections, but must be based on as much geochemical, petrographic and stratigraphic information as circumstances permit to be obtained.

K.H. WOLF AND G.V. CHILINGARIAN

78

mentary, volcanic, plutonic, and metamorphic rocks. The sedimentary and volcanic rocks may have undergone diagenesis before erosion and transportation and, therefore, show features that suggest the conditions of secondary processes in the source-rock environment. Factors controlling diagenesis Certain factors will initiate diagenesis, and the same or other factors will perpetuate the old and/or cause commencement of new diagenetic processes. The sediments have a tendency to adjust to new physical and chemical conditions and would, theoretically, reach equilibrium. The micro- and macroenvironmental conditions above and within the sediments, however, change continuously. Sometimes, equilibrium may be established, as, for example, in cases where unstable feldspar grains are completely replaced by clay minerals. In many cases, however, the physical and chemical conditions shift so rapidly that only a small fraction of the reaction involving the sedimentary framework reaches equilibrium. In particular during the early diagenetic stages numerous successive and overlapping processes will be acting at a relatively fast rate on both micro- and macroscales, when movements of interstitial fluids are at a maximum, biological activity is producing chemically active substances, maximum pore space is available, and temperature change is more or less sudden due to diurnal exposure. The following factors influence diagenesis of sediments: (1)geographic factors (e.g., climate + humidity + rainfall + type of terrestrial weathering surface water chemistry); (2) geotectonism (e.g., rate of erosion and accumulation, coastal morphology, emergence and subsidence, whether eugeosynclinal or miogeosynclinal); (3) geomorphologic position (e.g., basinal versus lagoonal sediments + current velocity + particle size sorting + flushing of sediments); (4) geochemical factors in a regional sense (e.g., supersaline versus marine water and volcanic fluids and gases); (5) rate of sediment accumulation (e.g., halmyrolysis ion transfer + preservation of organic matter + biochemical zonation); ( 6 ) initial composition of the sediments (e.g., aragonite versus high-Mg and low-Mg calcite and isotope and trace-element content); (7) grain size (e.g., content of organic matter -+ number of bacteria + rates of diffusion); (8)purity of the sediments (e.g., percentage of clay and organic matter -, base exchange of clays altering interstitial fluids); (9) accessibility of salldstone framework to surface (e.g., cavity systems permit replacements) ; -+

-+

-+

DIAGENESIS OF SANDSTONES AND COMPACTION

79

(10) interstitial fluids and gases (e.g., composition, rate of flow, and exchange of ions); (11)physicochemical conditions (e.g., pH, Eh, partial pressures of gases, and COz content); (12) previous diagenetic history of the sediment (e.g., previous expulsion of trace elements will determine subsequent diagenesis). The numerous large-scale environmental parameters listed above influence in one way or another the more local environments and these, in turn, influence the microenvironments. There is a complete gradation and overlap of these macro- and microfactors as one example below illustrates: climate + geomorphology particle size

4

amount and type of bacteria

4

rate of diagenesis

4

pH and Eh

4

type of replacement The actual processes that lead t o diagenetic alterations and modifications of sediments can be divided as follows. (1) Physicochemical processes: solution, corrosion, leaching, bleaching, oxidation, reduction, reprecipitation, inversion, recrystallization, cementation, decementation, authigenic mineral genesis, crystal enlargement, replacements, chemical internal sedimentation, and aggregation and accretion. ( 2 ) Biochemical and organic processes: accretion and aggregation, particlesize reduction, corrosion, corrasion, mixing of sediments, boring, burrowing, gas-bubbling, breaking down and synthesizing of organic and inorganic compounds. ( 3 ) Physical processes: compaction, dessication, shrinkage, penecontemporaneous internal deformation and corrasion, and mechanical internal sedimentation. Many of the above processes are commonly considered syngenetic. As they can occur within the sediments and directly alter and influence diagenesis, however, they must be considered as part of diagenesis. It is the total or collective influence of all factors that must be examined in a final analysis. As Krumbein (1942)pointed out, variations in the diagenetic end-products may occur either-with different sediments in the same environment, or with the same kind of sediment in different environments.

TABLE 3-111 Various lithogenetic stages as defined by different authors in English-Americanliterature (after Dunoyer de Segonzac, 1968, table 4, p. 189)

-

T h e detrttal particles still in movement in the water

1 -*

a

in a sediment wlth o high water content but isolated from the environment of sedimentation

T h e sediment has became a'more ar less compact rock

TWENHOFEI

PETTIJOHN

1926. 1939.195

(1949.1957)

I

1

I

WILLIAMS TURNER GILBERT (1954)

I

iI

,-on,,y

I

(1960)

j

T Depor,r,on

DlAGENESlS

-I-----I

~

I

~

~

~

.

.

I

I

METAMORPHISM

. . . .

I lrrh,frcol,on

'

I' . . . . . 1: . . .. . .. . . . __._.... ~ ~ ._._____.__ ~ _ .~ . ~ ._... . . .. . . . . . . . . The sedimentary . ... . . . . . . . series finds itself . ,. . ,. . . . .. . .... .. ... . , . .. .. . . . .. .. . under metomarphic . ., . . . . .. . .. . ... .. .. ... conditions on account ......... : ... . . . . .. .. .. . .. .. .. . ..... .. .. ... ...... . of orogeny . . .. .. ... .. .. . .. . ............... . . .. . .. . .. . . . ,

-

'

'

_ _ _ ~ ~ ~ ~

'

-

Tectonic phenanena place the sediments r r r r p y w under conditians of Xofomorph,sm decompression and 'del!lk,ficoIron, leaching. 01 exposed in outcrops

t

* In reality parts 11 and 111 represent very unequal thicknesses of sediments, tens of meters for part 11, hundreds or even thousands of meters for part 111.

DAPPLES (1959)

I

DAPPLES (1962)

DIAGENESIS OF SANDSTONES AND COMPACTION

81

As pointed out already, the lithogenetic stages used in the study of sedimentary rocks have been established differently by various investigators, an excellent summary of which has been provided by Dunoyer de Segonzac (1968). His summary tables, giving the stages of lithogenesis, are presented here in Tables 3-111, 34V, and 3-V, as based on the English, German, and Russian literature, respectively. Compaction is a diagenetic process and one may even speak of “compaction diagenesis”. This book is devoted mainly to this particular process, which may occur either as a physical and/or physicochemical and biochemical phenomenon. Compaction of sediments is the process of volume reduction expressed either as a percentage of the original voids present or of the original bulk volume. Although the process affects mainly loose, unlithified sediments, it may also have profound influences on well-cemented deposits, TABLE 3-IV Various lithogenetic stages as defined by different German researchers (after Dunoyer de Segonzac, 1968, table I, p. 159)

*

In reality parts IT a d 1 1 1 represent very unequal thicknesses of sediments:tens of meters for part 11, hundreds o r even thousands of meters for part 111.

TABLE 3-V Russian nomenclature related to diagenesis (after Dunoyer de Sgonzac, 1968, table 2, p. 168) - -. ~~~~~

AUTHORS

I F LITHOGENESIS

I

F ERS MAN j1922)

PUSTOVALOV

S H V ET SOV

TEODOROVICH

RU K H I N

(1933, 1 9 4 0 )

(1934, 1957)

(1961)

(1961 )

The detrital particles stlll in movtrnent in the water Particles immobilized in o sediment with o high water content but isolated from the environment of sedimentaiion

.*.

................

T -

The sediment hos become a more ar less comDact r o c k

SEDIMENTS

T h e sedimentary series finds i t s e l f under metamorphic conditions on accouni of orogeny

v*

Tectonic phenomena place the sediments under conditions of decompression and leoching. or exposed in outcrops

I n reality parts I 1 and 111 represent very unequal thicknesses of sediments:trns of meters for part 11, hundreds or even thousands of meters for part 111.

VA S SOEV I C ti (1962)

1

(1958,1963)

I

DIAGENESIS OF SANDSTONES AND COMPACTION

83

as will be discussed below. The intergranular spaces of clastic and detrital sediments are eliminated by closer packing, crushing, deformation, expulsion of fluids, and, possibly, dissolving of grains. Cemented sediments may undergo compaction through solution along stylolites, for example. Krumbein (1942) gave the following average values of porosities of freshly-deposited material: sand = 45%, silt = 50-65%, mud = 80-90%, and colloids (less than 1mu in diameter) = approximately 98%water. According to Miiller (1967, p. 135), the initial water content of argillaceous muds is approximately 50-80%, which corresponds to a porosity of 70-90%; whereas porosity for sands is only 30-5096, which corresponds to 20-30% water content. S e m y a (1969), however, found an average of about 150% water* (on dry weight basis) for the upper 200 cm of the lemanic sediments of Lake Geneva. She also reported that the water content of the very first layer of sediments was 250% $ when measured on an undisturbed sample. The degree of compaction generally depends partly on the ratio of fine to coarse material and on the character of the sediment framework. Inasmuch as mixtures of fine and coarse clastic grains are quite common, in this chapter a reference will be made occasionally t o clay-sized material, though the book is devoted to the compaction of coarse-grained sediments. One should note here, that the behavior of a mixture of different grain-size classes may or may not lie somewhere between the behavior of the separate grain-size classes, but little is known about this subject and more research work is required. Although compaction is part of diagenesis, if one makes compaction the “center of consideration”, then diagenesis could be subdivided into pre-, syn-, and postcompaction stages (Table 3-VI). On first sight this may look to be an artificial division; however, in detailed work on, for example, the relationship between fluid movements during diagenesis and their influence on the origin of various types of chemical cements, including metalliferous ores (see Chapter 5), such an approach should prove to be useful. Little is known about the diagenetic features formed by “water of compaction”, in contrast to those produced by other varieties of intrastratal fluids and mixtures of solutions of different derivation. Consequently, in future investigations some type of genetic classification of cements and other diagenetic features formed by different types of solutions and a corresponding nomenclature will become necessary. The various sources of chemicals that form cements in sandstones are presented on p. 76 above; the reader is also referred to Chapter 5 on ore genesis in sediments by compaction fluids. Compaction, as illustrated and defined in Tables 3-111, 3-IV, 3-V, and 3-VI, is not clearly defined by an upper and lower time- and/or space-bound*Weight of water divided by the weight of dry solids.

03

TABLE 3-VI General relationship of the process of compaction to diagenesis DIAGENESIS

lb

1 - -+

EPIGENESIS I

J

METAMORPHISM

+

Precompaction

_2

-

syncompaction -

*

noncompactional fluid movements in general (i.e., fluids of various origins)

postcompaction

-

-1

WEATHERING

-

e--

<

“water of compaction” movements

-

-chemical precipitation from all varieties of fluids

,

x

I=

chemical precipitation from compaction fluids

DIAGENESIS OF SANDSTONES AND COMPACTION

85

ary, so that such tables must be used with extreme caution. The rate of

compaction and the decrease in porosity and permeability, as well as the rate of expulsion of fluids, will change with time both vertically and horizontally in sedimentary basins. In a study of release of chemical elements, e.g., lead, into compaction fluids and their migration from the fine-grained into coarser-grained sediments, one would have t o consider the changes in the properties of the sediments during compaction with geologic time, and any “definition” of compaction-diagenesis and its paragenesis is to be taken as a guide rather than a rule. The amount, rate, and mechanism of release of fluids are controlled by: (1)changes in permeability, which also controls the rate of flow of fluids; (2)temperature increase during burial; and (3) dissociation of water during temperature and pressure increases with increasing depth of burial (Blatt et al., 1972, fig. 7.1).Several case histories are presented in the various sections of this chapter. The study of diagenesis has been mainly an observational science until relatively recently when geologists and geochemists have obtained data on natural and artificial (= laboratory) diagenetic, chemical systems. More quantitative data is available now on diagenetic processes than ever before. The interpretation of these data will be assisted by the development of new theories. Berner (1971, 1972), for example, presented theoretical models, e.g., chemical kinetic models for steady and non-steady diagenetic processes. He also quantified, or offered formulae for, total compaction, rate of compaction, rate of flow through a horizon, and total volume of water passing through the sediments. DIAGENESIS IN SANDSTONES

The literature on the various aspects related to sandstone diagenesis is voluminous (e.g., see Pettijohn et al., 1972) and for the purposes of the present chapter, the publications by Dapples (1962,1971,1972)are particularly useful. On the basis of mineral associations, intergrowth, and replacements, he recognized that four oxide series among the sandstones are not purely composed of silica (or quartz) but have other constituents: (1)alumina, lime-magnesia, iron oxide series; (2) silica, lime-magnesia, iron oxide series; (3) silica, alumina, iron oxides series; and (4) silica, alumina, limemagnesia series. Based on simple mineralogy, he listed (see Table 3-VII) the recurrence of certain types of mineral associations showing secondary growth and mutual replacement. On using the four oxide series, Dapples (1962, figs. 1 to 4, pp. 915-929) showed that chemical components are combined to form specific minerals, which he discussed in detail. One should note that in his diagram, Dapples listed minerals (sometimes end-member

86

K.H. WOLF AND G.V. CHILINGARIAN

TABLE 3-VII Minerals showing mutual intergrowth and replacement in poorly-cemented sandstones of simple mineralogy (after Dapples, 1962, table 1, p. 915) ( 1 ) Quartz-hematite

(2) Quartz-calcite (3) Quartz-hematite-calcite-siderite ( 4 ) Quartz-chert-clay mineral (5) Quartz-chert-clay mineralsiderite (6) Quartz-glauconite-clay mineral (7)Quartz-calcite-clay mineral (8) Quartz-glauconite-clay mineral-calcite

minerals with minerals of transitional composition in between) that are stable under near-surface, early-burial, and late-burial conditions. Dapples (1962) explained that according t o the type of chemical reactions and the major sequence of occurrence, three diagenetic stages are recognizable (Table 3-VIII): Stage l-redoxomorphic, which comprises the episodes of sediment accumulation and early burial. The principal chemical reactions are reduction and oxidation. Stage 2-locomorphic, involving the replacement of one mineral by another without the minerals entering into mutual reactions, which is typical of the early-burial stage and forms an important part of the process of lithification. Stage 3-phyllomorphic, occurs subsequent to the locomorphic replacements and involves the origin of micas, mainly as a transformation product from clay minerals. In sandstone petrology, the progress of diagenesis may be considered t o be terminated at either the locomorphic or phyllomorphic stage. It must be recognized, however, that the phyllomorphic stage overlaps with the zeolite and chlorite stages of burial or low-grade metamorphism. Many other investigators have discussed the problem of gradational change from diagenesis into metamorphism. Blatt (1966) mentioned that, according to metamorphic petrologists, diagenesis per se stops short of the zeolite facies in tuffaceous sandstones and before recrystallization of quartz in orthoquartzites, arkoses and lithic sandstones. These limits imply temperatures of approximately 10°C to 200 f 50°C and pressures of one bar to about 2000 bars. Pettijohn et al. (1972) used the term diagenesis in a wider sense, namely, that it affects the sediments up to the lowest grade of metamorphism (the greenschist facies). There is no definite demarcation between diagenesis and metamorphism. As a sandstone is buried deeper and deeper in a sedimentary basin and heat and pressure increases, a stage will be reached where one can call the rock either sedimentary or metamorphic depending on the classification and nomenclature adopted. Both diagenetic and metamorphic stages

DIAGENESIS OF SANDSTONES AND COMPACTION

87

TABLE 3-VIII Reactions characterizing diagenesis (after Dapples, 1962, table 2, p. 931) ~

Redoxomorphic stage (reversible reactions)

~~~

Locomorphic stage (replace- Phyllomorphic stage (unidirecment reactions) tional reactions)

+ muscovite

aragonite. by calcite

“clay minerals” or biotite

Hematite + calcite T?.siderite

calcite by dolomite

montmorillonite -+ chlorite

Hematite + calcite + Mg”’ siderite + ferrodolomite + dolomite

carbonates by quartz or chert

“clay minerals”

*

-+

chlorite

Hematite + clay minerals + quartz or chert by carbonates “clay minerals” + Fe+* silica + chlorite + greenatite lite + stilpnomelane (?)

-+

bio-

Hematite + chlorite + chamo- feldspar by carbonates site (?)

“clay minerals” + chert rite

Hematite + illite + glauconite

quartz, chert, clay minerals by carbonates

“clay minerals” + quartz -+ sericite

“Bauxite” + silica nite

opal by chert or quartz

kaolinite + glauconite or. illite + Mg+2 -+ chlorite +

silica solution =+chert or quartz

glauconite

+ kaoli-

K+

-+

-+

chlo-

feldspar

Diaspore or boehmite + silica + clay minerals

illite or glauconite + muscovite

Diaspore (?) + silica + K+ + clay minerals (glauconite)

kaolinite + illite + glauconite + calcite + Mg+2 -+ muscovite + biotite + feldspar + dolomite + chlorite

Bauxite + hematite + silica (minor) =+clay minerals + pyrite

illite or glauconite + calcite + Mg+2 + micas + feldspar + dolomite

Kaolinite + K+f i illite

feldspar

Biotite

plagioclase -+ chlorite + chert

* glauconite

Feldspar chert

=+clay minerals +

Glass montmorillonite + chert -+

-+

-+

chlorite

sericite

88

K.H. WOLF AND G.V. CHILINGARIAN

leave imprints on rocks, e.g., on graywackes. As Pettijohn et al. indicated, although there is a diagenesis-metamorphism continuum in physical variables, e.g., temperature and pressure, there are important differences between the two stages. Metamorphic petrology involves the genetic interpretations of secondarily formed mineral assemblages, controlled by bulk chemical composition, as indicators of pressures and temperatures, assuming equilibrium or a close approach to it. Very often there is no original sedimentary feature left, except under certain circumstances, e-g., in the case of burialmetamorphosed graywackes. This approach results in an order of mineral assemblages in the metamorphic rocks expressed either by isogrades, metamorphic facies or petrogenetic grids. On the other hand, the sandstones, even after diagenesis, are composed of mineral assemblages that tend to reflect the composition of the original, obviously non-equilibrium detrital mixtures more than the effects of pressure and temperature, because the maximum values of the latter two variables involved in diagenesis are much lower than in metamorphism. In addition, the rates of reactions of mineral neomorphism are slow at low temperatures. Certain equilibrium assemblages of very low-grade burial metamorphism (considered late diagenetic or catagenetic stage by some investigators) may be the result of changing composition of intrastratal fluids rather than the increase in temperature and pressure, e.g., in pyrodastic rocks. Pettijohn et al. (1972) offered a diagrammatic representation of six stages t o which a sandstone is exposed during burial (Fig. 3-7). This type of scheme is particularly useful, because they attempted to assign clay-mineral reactions for each stage as shown in Table 3-IX. For a good summary of sandstone diagenesis the readers are also referred to an excellent book by Pettijohn (1972, pp. 383-437), who divided the evidence of diagenesis into: (1) textural, (2) mineralogical, (3) physical, and (4) chemical. As shown later in separate sections, compaction and/or compaction fluids are involved in these diagenetic modifications. Pettijohn et al. (1972) also discussed the composition of sandstones in some detail. They stated that the silica minerals in sedimentary rocks are either clastic or of a secondary chemical and/or biochemical origin. The secondary chemical and biochemical processes can also alter clastic silica grains by dissolving, corroding and etching them. As conceptually shown in Fig. 3-8, the chemically formed silica is represented by the following minerals: opal, chert, or quartz (see also Carozzi, 1960, pp. 291-292, among others, for details). It is important to know that the source of S i 0 2 is fourfold: biogenic silica, volcanic glass, various silicates, pore waters, and detrital quartz grains during pressure solution. All these may release SiOz into compaction fluids which, in turn, may precipitate silica (and other minerals) when exposed to a different chemical milieu. Similar considerations apply to the origin of authigenic feldspars in sandstones which require

DIAGENESIS OF SANDSTONES AND COMPACTION Depth of burial

89

Microscopic Appearance Immediately after deposition. Ex& Original detritus. hieh porosity.

to air or water

of depo%itmnalenvironment

Buried a few meters to tens of meters. Exposed to Interstitialwaters Some wmpaclion. some wrly chemical precipitates possible.

Buried to modcrate depths of about ID00 m. Pore water may be a brine Chemical Cements may reduce porosity. days may be altered.

5 q burial to thousands of meters perhaps with folding. Porosity may bc "cry low

from chemical osminl and pressure solution.

Incipient metamorphism. Growth of chlorite and other metamorphic minerals wnlh extensive pressure solution and quaNitic texture.

@in and erosion. within tens of meters ol land surface. Invasion by meteoric watei. dcccmcntation and "wuthcrin8" of clays may increasz porosity. After

Fig. 3-7. The stages of diagenesis in relation to depth of burial and increase of pressure and temperature. (After Pettijohn et al., 1972, fig. 10-1, p. 387; courtesy Springer, New York.)

a suitable solution, possibly mobilized by compaction, to form overgrowths of albite or K-feldspar on clastic feldspar grains, for example (Fig. 3-9). Chemical conditions for the precipitation of feldspar are determined by the pH and relative amounts of various components in solution, i.e., K+,Na+, Mg2+, Ca2+, and Si02. The alkali metal/hydrogen ion ratio, at a minimum amount of SiO14- in solution, is indicative of the stability of feldspar with respect to a solution and thus of the possibility of its precipitation (Garrels and Christ, 1965, pp. 359-363; Pettijohn et al., 1972, p. 38). Hemley and Jones (1964)suggested that slightly elevated temperatures, that are associated with moderate to deep burial, are important in precipitating feldspar. Diagenetic processes are not uniform and regular as demonstrated by sandstones that are hundreds of millions of years old and only slightly cemented in contrast to more recent, well-cemented sediments. Also, several sandstones with the same degree of lithification may have had very different post-lithification * histories. In general, during increasing alteration of different types of sandstones, there is an eventual convergence to a chemical

K.H. WOLF AND G.V. CHILINGARIAN

90 TABLE 3-IX

Some clay-mineral reactions during sandstone diagenesis (after Pettijohn et al., 1972, fig. 10-2, p. 431) Clay mineral formed

Precursor

Components added to (+) or subtracted from (-)

Stages of diagenesis (see Fig. 3-7)

Kaolinite

feldspar

1,2,6

Kaolinite Illite

pore space kaolinite

Muscovite

kaolinite

Illite

montmorillonite

--(K+,Si02) +HzO* +(A1203, SiOz, H2O) +(K+, Si02) -(A1203, HzO) +K+ -Hz 0

Chlorite

montmorillonite

Montmorillonite

volcanic glass

Glauconite

illite

+K+

-(SiOz, HzO, Na+, Ca2+, Mg2+, Fe2+, etc.) +(Fez+, Mg2+) -(SiOz, HzO, Na+, Ca2+)

+HzO

-(Na+, K+, Ca2+) +(Fez+, Fe3+) -(K+, A1203 1

* 2KAISi308 + 2H+ + 2HCOT + 9 H 2 0

+

2,6 3,4,5 5

3,4 3,4,5 1,2,3,4 192

AlzSizOs(OH)4 + 4H4Si04 + 2K+ + 2HC03.

equilibrium in both mineralogic composition and texture (as long as the original bulk chemical composition is not very different), but certain initial differences can persist into the higher grades of alterations. A fresh sand is a porous, non-equilibrium assemblage of clasts the composition of which is I

Pressure SOlUllO"

Volcan,c glass and other silicate\

iJIACIENF5IS

PROVFYAhCF

DIAGENESIS OF SANDSTONES AND COMPACTION

91

1

Fig. 3-9. The origin o f feldspar in sandstones. (After Pettijohn et al., 1972, fig. 2-4, p. 38; courtesy Springer, New York.)

“derived” from the source areas. During diagenesis, that includes compaction, cementation, and other alterations of various types, there is a loss of unstable detrital grains and an increase in stable authigenic components. After long and deep burial a quartzitic arenite would become a well-cemented quartzite (= quartz clasts cemented by quartz cement). Lithic arenites will show complete cementation by a combination of quartz, carbonate and clay minerals, the latter commonly represented by an illite-chlorite assemblage. (This complex material makes it difficult to draw a definite boundary between “cement” and “matrix”, as discussed below.) The final products of all the secondary alterations of sandstones reflect the temperature and pressure increases, composition of pore fluid chemistry, original composition of the sediment and its texture, and geologic time. The total post-depositional geologic history of the sediment was involved. Fiichtbauer and Muller (1969) presented the textural and mineralogical results of sandstone diagenesis (Table 3-X), including the influence of compaction. Each one of the diagenetic processes is treated at some length in subsequent sections of this book. Practical applications of detailed diagenetic and related studies of sandstones are exemplified by the work of Griffiths (1964, and his numerous other publications). Some examples are provided in this chapter. Griffiths (1964, p. 640) investigated several thousand core samples of barren and oil-producing sedimentary rocks, and showed that the latter exhibited specific characters which differed in degree from those of the oil-bearing but non-producing sands. Both producing and non-producing oil-bearing sands,

K.H. WOLF AND G.V. CHILINGARIAN

92 TABLE 3-X

Textural and chemical-mineralogical results of sandstone diagenesis (free translation of table 3-14,p. 105 in Fuchtbauer and Muller, 1970) Chemical process

Example

Influence on: porosity compaction

Dissolution (a) pressure-solution (b) intrastratal solution

pressure-quartzite dissolved heavy minerals

(-1 strongs ( + ) filledS

strong none

( + ) strong

none

Neoformation (partly cementation) (a) autochthonous1 sand grains grew at the expense of siliceous matrix homogene2 quartz grains dissolved at pressure points; SiOz precipitated in pressure shadows (b) autochthonous kaolinite formation in pores heterogene3 from feldspars, which were dissolved in other parts of the sandstone (c) a l l o c h t h ~ n o u s ~ silicification as a result of SiOz homogene supply from an external source cementation by anhydrite (d) allochthonous heterogene

,

Replacement (a) autochthonous (b) allochthonous

* Autochthonous

kaolinization of feldspar dolomitization of orthoclase

(-) doubled

( + ) strong

none

(-) filled

none

(-) filled

none

( + ) strong none

none none

= derived from the components of the sandstone (= internal source); homogene = neoformation of a mineral already present; heterogene = neoformation of a type of mineral not previously present in the sandstone; allochthonous = delivered from an outside source; ti (-) = reduction in porosity and (+) = increase in porosity

*

in turn differed from the barren sediments. Griffiths stated that this is an example of different textural properties of the sediments showing complex interdependent relationships (see also Griffiths, 1961). According to him (1964, p. 640), laboratory experiments have shown that sands having different grain sizes, exhibit different saturation characteristics (Griffiths, 1957):

DIAGENESIS OF SANDSTONES AND COMPACTION

93

“when a fluid-saturated sand is invaded by a second fluid, the second fluid is preferentially concentrated in the coarse-grained sands. For example, in the secondary uranium ores on the Colorado Plateau, when differential saturation occurs, the ore is selectively confined to the finer-grained sands. Since the ‘second fluid’ is expected to saturate the coarse-grained layers, this suggests that uranium saturation was original and the second fluid removed the ore from the coarser-grained sands.” Uranium concentrations have been found in coarser silts associated with very fine sands in the Entrada Formation, whereas they occur in very fine sands associated with fine to medium sands in the Salt Wash Member of the Morrison Formation. Uranium is also concentrated in fine-grained gravel to coarse-grained sands associated with coarser gravels in the Shinarump deposits. (Jobin, 1962, presented similar explanations for this distribution pattern.) In some cases reported by Griffiths (1964,p. 641), carbonate cement occurs in the coarser-grained sandstones, showing that the carbonate-bearing solutions invaded the sedimentary pile displacing earlier fluids. Griffiths stated: “It may well be argued that the controlling factor is not really grain size but some property which varies in a similar manner; however, since the variation in grain-size is closely associated with that in the controlling factor, grain-size variation is a convenient and efficient means of following the critical changes which are associated with, and in some complex way control, the differential saturation.” Of course, other variables are influential in addition to grain size, e.g., amounts and compositions of matrix and cement, sorting, orientation, packing, and degree of compaction, in determining porosity and permeability. But the variations in the several parameters, many of which are related to grain size, are interrelated and interdependent. As Griffiths (1964,p. 642) suggested, the question is how much additional effect, to that supplied by changes in grain size, is supplied by variation in the type and amount of cement and/or matrix and several other textural parameters. All of them should be considered step-by-step to determine their influences on porosity and permeability. Compaction, of course, which is our main interest in this book, will have to be considered here also. In order to evaluate the influences of several petrographic parameters on the mass properties of sediments, Griffiths (1964,p. 642) proposed to use statistical methods to determine: (a) the most important controlling property; (b) how much control this property exercises over the mass properties; (c) what additional information on control is supplied by adding new properties; and (d) what is the most parsimonious combination of properties which accounts for the greatest degree of control over porosity and permeability, for example. The degree of predictability and the interrelationships are determined by means of multiple regression and by using discriminant equations. The latter method was employed by Griffiths (1964,p. 649) in the

94

K.H. WOLF AND G.V. CHILINGARIAN

TABLE 3-XI Comparative measures of interrelationships between porosity and petrographic variables in some Berea oil-producing sands and in some barren Pocono sandstones (after Griffiths, 1964, table 1, p. 644) Variable

Degrees of freedom

~~

Berea

Pocono

0.7647*** 0.0265 0.15 08 * 0.0034 0.0011 0.16 31 0.1833*** 0.0149 0.08226

0.0062 0.4700 0.0055 0.4141 0.3995 0.4393 0.2891 0.1586 0.2108

Percent variation explainedl Berea

Pocono

62.0 2.2 12.2 0.3 0.1 12.4 89.2 10.8 100.0

0.2 14.9 0.2 13.1 12.6 13.9 54.9 .45.1 100.0

~

Packing Additional for grain size % Quartz Orientation Size sorting Shape Cumulative total Residual Total 1

Mean square

1 1 1 1 1 1 6 9 15

*

Percent variation = Sxf/Sx,2,t.

example given in Table 3-XI. Samples from the Berea oil-producing sandstone were compared with those of the barren Pocono sandstones, using the following petrographic properties: (1)mean proportion of quartz; (2) grain size (Le., the average length of the long axis of quartz grains); (3) size sorting (i.e., the standard deviation of long axis of quartz grains); (4) shape (i.e., the ratio of short over long axis of the quartz grains); (5) perfection of preferred orientation (i.e., standard deviation of axial inclination in degrees); and (6) packing of the quartz grains. Upon testing, Griffiths (1964, p. 643) was able to reduce the discriminant equation based on parameters 1 to 6 to an equation where packing, shape, and grain size led to maximum discrimination. Thus, the evaluation of these three parameters is adequate to differentiate the oil-bearing sandstones from the barren sandstones. The porosities differed in these two sets of samples and separate multiple regressions showed that the relationships between porosities and the petrographic variables also differed in the two cases: the porosity and petrographic variables of the Berea specimens showed a high degree of predictability, which was absent in the case of the Pocono sandstones (Table 3-XI). The data of the Berea sandstone indicated that the packing, proportion bf quartz, and grain shape “explains” or “accounts” for 88.9% of the variation in porosity, whereas in the Pocono specimens all above-mentioned six properties account for 54.9% variation. It is particularly interesting to note that no single variable controlled significantly the predictability of the porosity. Together with other

DIAGENESIS OF SANDSTONES AND COMPACTION

95

studies, it is evident that certain properties, in this case porosity, is controlled by other different properties in different sandstones. The above is a particular case of investigation, but Griffiths (1964) went beyond it to find a more general model, which is of special relevance to the present theme of the chapter. He pointed out (p. 647) that from studies of the petrology of detrital sediments, the dominant factors are source area, processes of sedimentation, and diagenesis. According t o him, a model can be established in which these three factors may be related to the measured properties of the sediment. This model can be used to compare roles played by different properties under varying genetic circumstances. Table 3-XI1 is an idealized model based on first approximation in which, for example, the percentage of clastic components is related t o source area and weathered source area materials, but is independent of other factors and properties. In reality, however, the simple relationships presented in Table 3-XI1 are more likely to be more complex, as shown in Table 3-XIII. In this table, the useful property of independence is largely lacking and, as Griffiths pointed out: “Effects of subsequent factors modify the relationships of properties to factors and may completely subdue the original relationships. Interdependencies among the properties are thus the rule rather than the exception. It would be useful, therefore, to find an analytical tool which could in some way separate out the various effects and, where appropriate, introduce indeTABLE 3-XI1 Idealized component analysis reflecting relationships among provenance, genetic process and properties of sediments (after Griffiths, 1964, table 2, p. 648) Property

Model character (C)

C1 source area Proportion of: Detritus Matrix Cement

c3

c2

weathering

xxxx xxxx

erosion

transportation

deposition

diagenesis

0 0 0

0 0 0

0 0 0

xxxx xxxx

0

0 0

0 0

Size Sorting

0 0

0 0

xxx

xxx 0

xxx xxxx

0 0

Orientation Packing

0 0 ‘

0 0

0 0

0 0

xxx xxx

xxx xxx

0

K.H.WOLF AND G.V. CHILINGARIAN

96

TABLE 3-XI11 Relationships among provenance, genetic process and properties of sediments (after Griffiths, 1964, table 3, p. 649) Property

Model character (C)

source area

c3

C1

c2

erosion

transportation

deposition

xxxx xxx xx xxx X

X X 0

X

X

X X 0

Size Sorting

X X

X X

xx xx

xxx xxx

Orientation Packing

0 0

0 0

0 0

X 0

xxxx xxxx xxxx xxx

Proportion of: Detritus Matrix Cement

weathering

xx X

diagenesis

X

xxx xxx x x xx xxx

pendence, so that the effects of the various factors on the variation in the measured properties may be separately evaluated.” Component analysis, which is one technique among many in factor analysis, is such a tool (Griffiths, 1962). In its use, the parameters are reduced to three for simplicity, i.e., C1,C2,and C3 in Tables 3-XI1 and 3-XIII. An example is given in Tables 3-XIV, 3-XV, and 3-XVI. The matrix of coefficients of correlation of zeroorder are given in Table 3-XIV for eleven properties of a Mississippian quartzose, low-rank graywacke (see Pettijohn’s classification in the Introduction chapter) of the Maxton Sandstone. Table 3-XIV indicates that there are: (a) a typical high correlation ( r > 0.95) between long “a” and “b” axes of quartz grains; (b) moderately high degree of association between the standard deviations (8, and 6,) of these axes, i.e., among size sorting of these axes; and (c) association between porosity and the log of permeability ( r = 0.85). Apart from these features, it is difficult to make an interpretation of the other data because of complex interdependencies present among the measured variables. When the data is treated by component analysis, the characteristic roots (last row in Table 3-XV) demonstrate that five components account for some 90% of the variation, i.e., the relationships among the eleven variables can be represented in terms of five components. As shown in Table 3-XV, the first component (C,)“explains” 50.6% of the variation in the relationships among the variables, the second component (C,) an

TABLE 3-XIV Matrix of correlation coefficients for eleven properties of the Maxton Sandstone, Mississippian, West Virginia (after Griffiths, 1964, table 4, p. 652) Mineral

Quartz Rock fragments Matrix Silica Xa

zb

aa &b

8

Quartz Rock frag- Matrix ments

i

Silica

Grain size

1

Size sorting

(%a

(jZb

(6,)

($1

Orientation (6")

Porosity

Log. permeability

~

-0.230

-0.441

-0.256

-0.186

-0.147

1

-0.512 1

-0.505 0.382 1

-0.580 0.811 0.342 1

0.581 0.781 -0.336 0.993 1

0.048

0.049

0.555 -0.347 -0.322 -0.263 -0.276 1

0.427 -0.274 -0.345 -0.255 -0.280 0.724 1

Porosity Log. permeability For n = 33, rij 2 0.344 significant at Po5 level; subscripts a and b refer to long and short axis, respectively.

-0.225

-0.044

0.534 -0.471 -0.190 -0.558 -0.589 0.241 0.172 1

0.702 -0.667 -0.606 -0.709 -0.700 0.416 0.185 0.477 1

0.096 0.575 -0.617 -0.678 -0.617 -0.605 0.386 0.217 0.306 0.850 1

98

K.H. WOLF AND G.V. CHILINGARIAN

TABLE 3-XV Component matrix of Maxton sandstones (after Griffiths, 1964, table 5, p. 654) Component

C1

Variable Quartz Rock fragments Matrix Silica Grain size a Grain size b Sorting a Sorting b Orientation Porosity Permeability Variance accounted for

103(rij‘) = relative loading -504 784 145 351 -229 7 95 341 -140 -8 29 -095 -379 -631 325 147 -867 287 177 -865 637 285 561 645 346 461 0 28 -554 608 -076 -076 880 -071 098 814

c3

c2

50.6

64.4

77.3

c4

c6

213 -119 -242 5 20 -232 -246 205 364 216 -321 -436

145 -01 5 022 -369 111 089 -161 062 458 -1 34 -184

86.5

90.8

additional 13.8%,the third component (C,) another additional 12.9%, and so forth. All components beyond the fifth (i.e., 6 to 11 or C6 to C l l ) are considered to represent random variation totalling some 9%. To attempt an interpretation of this data as based on the idealized models of Tables 3-XI1 and 3-XIII, Griffiths (1964) offered the simplified Table 3-XVI. According to him (p. 655): “The largest loading in the first component is average grain size and so this component is considered to represent effects of the factors erosion, transportation, and deposition. Since both proportion of rock-fragmen& and matrix show their heaviest loadings on this component also, it is deduced that variation in these detrital constituents is largely affected by selective sorting on the basis of size. If this interpretation is correct, then the matrix material is largely detrital in origin.” Griffiths pointed out further that inasmuch as porosity and permeability also show their heaviest loadings on this component, then these dependent properties are most closely associated with the average grain size. Thus, the porosity and permeability of this reservoir rock is largely dependent on the same factors that determined the size distribution, and the other properties are affected to some degree by the effects of the factors dominating this component. Griffiths also showed that the second component exhibits its heaviest loading on the two measures of size sorting. This component is independent of (and orthogonal to) the first component and, consequently, this may be interpreted as a subsidiary but important effect of selective sorting, i.e., the first two components both reflect selective sorting on the basis of size and shape, the one representing,

DIAGENESIS OF SANDSTONES AND COMPACTION

99

TABLE XVI Simplified component matrix of Maxton sandstones (after Griffiths, 1964, table 6, p. 656) Component +

1

4

5

X

+

X

+ +

+

2

3

xx

+++

Variable Quartz

+ +

+++

Rock fragments

xxx xx xxxx xxxx

Matrix Silica Size Z, fb

Porosity

++ + ++ ++++

Permeability

+++

Sorting ab

Orientation

+ + ++ ++

xx

+ X X

Explanation: (rij’)1o3

>850 700-850 500-700 250-500 <250

positive

++++

+++ ++ +

blank

negative

xxxx xxx xx X

for example, the average current velocity and the other the range in current velocities. Griffiths (p. 657) stated that ‘‘. . . the proportion of quartz possesses a heavy loading on this second component and so a large part of the variation in quartz proportion is a reflection of control by factors affecting fluctuations in current velocity. Again grain size and proportion of rock fragments and matrix are affected to some degree by the component. Since some two-thirds of the total variation (64.4%) is accounted for by these first two components, the remaining components are not nearly so distinct. It seems likely that as proportion of silica cement is dominant in the 4th component and present in the 3rd and 5th, diagenesis is the factor reflected in these components. The 3rd possesses its heaviest loading on quartz with a strong second on perfection of orientation, which suggests that the factors involved are processes operating at a late stage in deposition or early (physical?)

100

K.H. WOLF AND G.V. CHILINGARIAN

diagenesis. Porosity and permeability are also affected by the factors present in the 4th component, or essentially by diagenetic processes connected with the distribution of silica cements.” The reasons for presenting the above information from Griffiths’ publication is: (a) to demonstrate the complex interrelationships of numerous factors, which are discussed in this and other chapters; (b) to show the necessity of using statistical techniques; and (c) to point out that compaction will also have to be given consideration in the future investigations that range over thick vertical stratigraphic sections. The amount of matrix and cement and the degree of orientation and packing of grains, for example, may be controlled by compaction and interactions of sediments with compaction fluids. One useful approach may be that of combining Griffiths’ techniques with detailed statistical textural and fabric studies, e.g., determination of types of grain contacts (see pp. 133-163). CEMENTATION OF SANDSTONES CONTROLLED BY COMPACTION FLUIDS

The chemical precipitation of cements within the framework of sedimen-

tary deposits is dependent on a supply of chemical elements by intrastratal

solutions, usually, moving solution. Although it is conceded that compaction fluids are not the only types of fluids that can cause cementation, the processes of compaction and lithification must be fully comprehended in the investigation of diagenesis in general, in order to be able to determine not only the direct influences of compaction on cementation, but also to establish, for example, at what stage of compaction the sedimentary deposit was lithified. To provide some of the fundamental data on cementation processes, the following paragraphs will discuss a number of aspects. Although no direct reference may be made to compaction itself, it is understood that all the solutions thought to have been responsible for cementation in the cases treated, could be considered to have originated during compaction of a sedimentary unit under natural conditions. At the same time, all other conditions, such as an increase in pressure and temperature for example, can be visualized to be the consequence of burial in a sedimentary basin. One may justifiably inquire about the relationship between the amount of cement, porosity, and permeability, on one hand, and compaction, on the other. Although at the present time quantitative answers are not available, it is clear that the relationships between the amount of cement and porosity and between porosity and permeability are a function of the degree of compaction of the sediment. The stage of compaction when cementation occurred and the amount of porosity reduction as a result of compaction should be considered.

DIAGENESIS OF SANDSTONES AND COMPACTION

101

In a publication on carbonate cements in quartzose sandstones, Dapples (1971) listed two main groups, namely, (1)cement that did not cause destruction of the grain-supported framework, and (2) cement that did (Table 3-XVII). This subdivision indicates that in detailed studies of compaction features in sandstones it is of fundamental importance t o recognize the various textural relationships between the grains and the cement (and matrix, if present), because the time of cementation and the degree of compaction are some of the factors that control the fabric relationships between the framework and the secondary minerals. In particular one should note that Dapples’ group 11-A (Table 3-XVII) incorporates four varieties of cements that result in volume expansion during crystallization (see Spry, 1969, pp. 149-152). Such expansion has been noticed also by Wolf and Ellison (1971) in Pliocene volcanic arenites and rudites which were never covered by a thick overburden. It seems that this expansion as a result of precipitation of cement can only take place during early diagenesis, as long as the sedimentary unit is near the surface, i.e., when the overburden is relatively thin and compaction is at a minimum. Experimental work may eventually determine the exact values of pressures involved in crystallization and the maximum depth of burial at which expansion as a result of cementation can occur. Griffiths (1958) successfully attempted to relate porosity to petrography TABLE 3-XVII Physical classification of carbonate cement in quartzose sandstone (after Dapples, 1971, table 1,p. 197)

I. Cementation without destruction of the grain-supported framework: A. crystallization as a single event: (1) filling interstitial space as a simple adhesive: (a) without mineralogic reaction; (b) incorporating clay minerals already present in intergranular space B. crystallization as part of dual events: (1)filling pore space following partial welding of quartz grains ( 2 ) recrystallization into single or compound crystals following welding of cement to clastic carbonate grains in a mixture of quartz and carbonate grains

11. Cementation with destruction of the grain-supported framework:

A. by crystallization and volume expansion: (1)in interstitial pores to produce local expansion (2) in porous grains, causing rupture and rotation of such grains (3) around sand nuclei to form small concretions in situ ( 4 ) as large poikilitic crystals B. by replacement of silicate-mineral grains: (1)without destruction of bedding (2) by developmint of large concretions

K.H. WOLF AND G.V. CHILINGARIAN

102

t POROSITY, %

Fig. 3-10. Permeability (maximum) versus porosity in Cow Run Sand, West Virginia. (After Griffiths, 1958, fig. 1, p. 16; courtesy4 Sed. Petrol.)

and has found that in many oil sands porosity and permeability are related exponentially, so that when the latter is plotted on a log scale and porosity on an arithmetic scale, a linear relationship emerges (e.g., Fig. 3-10). Griffiths then concluded that as a first approximation any relationship between porosity and petrography also applies to the log of the permeability. Inasmuch as these interdependencies are not exact, they cause a scatter of the data around the trend line, as in Fig. 3-10, for instance. One should note that the relation breaks down in the 0-10% porosity and less than 2 millidarcies permeability ranges. As demonstrated in Figs. 3-llA,B, carbonate cementation caused the reduction in porosity, so that the degree of cementation would appear to be one of the most important petrographic parameters to consider. According to Griffiths (1958), there is an inverse relationship (approximately linear) between the porosity and the carbonate content (determined by employing an alkalimeter). It is also independent of geographic and stratigraphic location in this particular study; however, this need not be so in other cases, because the percentage of cement, as well as its type, often were directly controlled by regional variations in the sedimentary and diagenetic milieu. Below 5%, the porosity varies independently of the carbonate cement con-

DIAGENESIS OF SANDSTONES AND COMPACTION

I

1

1

1

x- Calculated o = Observed

,

103

E

Fig. 3-11. Relationship between porosity (Y)and percentage of carbonate cement ( X ) for Cow Run Sandstone, West Virginia. ( A after Griffitths, 1958, fig. 8 ; B after Griffiths, 1967a, fig. 21-5; courtesy McGraw-Hill, New York.)

tent and, in these instances, grain size, size sorting, and amount of silica cement are the main factors controlling the porosity of the sandstones. The maximum porosity of this sandstone is around 25%and when the carbonate

104

K.H. WOLF AND G.V. CHILINGARIAN

content exceeds this amount, there is no obvious relationship between carbonate content and porosity. The data discussed above was obtained from measurements on specimens from the same rock formation, which most likely have the same compactional history. Future investigations should include attempts to establish correlation among the degree of compaction, depth of burial, porosity (including “minus-cement” porosity; see Pettijohn et al., 1972,p. 424), and the amount of cement. Several examples of the regional distribution of cements in sandstones are presented here, which reflect upon the present state of knowledge. Mellon (1964)reported that discriminatory analysis of calcite- and silicacemented phases on the basis of their compositional and textural properties indicated that grain packing and/or pore size were the only variables associated with the distribution of cements. In his study, calcite-cemented phases were more loosely packed than adjacent silicate-cemented phases (i.e., chlorite, illite, quartz), but grain size, grain orientation, and detrital composition showed no significant association with the cement distribution independent of packing. Calcite-cemented sandstones tended to be coarser-grained on the average than silicate-cemented ones because of the marked correlation between pore size and grain size. The chlorite was formed first, entirely enveloping the more tightly packed constituents of the sandstone. Illite and quartz genesis subsequently filled the remaining parts of the larger pores. The passage of later carbonate-bearing fluids w a s controlled predominantly by the distribution pattern of the earlier cements. Of these, the calcite was preferentially precipitated in the more permeable, originally loosely packed framework of the sandstone, at the same time replacing some of the detrital and authigenic components. Thus, it appears that compaction has had an indirect influence on the distribution of the diagenetic cements, because it controlled the degree of grain packing and the porosity prior to the chemical precipitation of the calcite. But other factors also have to be considered. Garrison et al. (1969)showed that the composition of neither the Fraser River water nor the sea water of the same area could account for the precipitation of the early diagenetic carbonate cement in sandy sediments they investigated. They believed that the dissolution of carbonate shells by subsurface fluids and the precipitation higher in the stratigraphic section from the expressed water of compaction could have been responsible. As Pettijohn et al. (1972)mentioned, calcite cement replacing secondary quartz must be a later diagenetic product related to the redistribution of calcium carbonate within the sedimentary pile. The solubility of carbonates decreases with increasing temperature and increases with greater pressure, although the latter has a much smaller effect. Thus, the net effect of burial is to decrease the solubility of carbonates. This could account for some of the cement, but

DIAGENESIS OF SANDSTONES AND COMPACTION

105

unless large quantities of solutions are passed through the system, only small amounts of secondary carbonates would be produced. Burial results in an increase in pressure-solution, an explanation usually offered for the solution-deposition sequence of silica cement, but which can also be advocated for carbonate cementation. According to Pettijohn et al., the hydrostatic pressure effect on the solubility of carbonates is greater than that of quartz, which is reflected in the common occurrence of late carbonate cement preceding quartz cementation. One may mention here also that in coarse clastic units, higher salinity and higher temperatures and pressures favor the formation of anhydrite over gypsum cement. No data seems to be available on the direct control of compaction on the composition of evaporite cements. Heald et al. (1962) presented an interesting approach in their study of the origin of interstitial porosity in an ancient sandstone by testing the hypothesis, advanced by many investigators, that porosity is often of a secondary type as a result of leaching of carbonate (i.e., decementation) below an unconformity. The small calcite cement patches at corners in pores, observed by Heald et al., were believed to have been vestigial, i.e., residue left after dissolution. Evidence, however, was presented against the leaching concept, based on the presence next to the pore walls of quartz with unblemished secondary faces. Adjacent to calcite, the faces were imperfectly developed, so that one can conclude that leaching of this carbonate would have left voids lined by blemished quartz faces. Heald et al. gave several reasons to support their argument as stated above, including some based on laboratory experimental data. They treated the sandstone samples with acid to dissolve the carbonate before examining them to determine the proportion of the surface of each quartz grain covered by unblemished faces. The results are shown in Figs. 3-12 and 3-13,using a “crystal face index” from 0 to 10. 1

10,

CRYSTAL FACE INDEX (after artificial leaching)

Fig. 3-12. Relation between natural porosity and crystal face index after artificial leaching. I = Eastern outcrop area; 2 = Kanawha Co. subsurface samples; 3 = Tucker Co. subsurface sample. (After Heald et al., 1962, fig. 4, p. 294; courtesy J. Sed. Petrol.)

K.H. WOLF AND G.V. CHILINGARIAN

106

0

' 8

I 6

4

2

CARBONATE. % (before artificial leaching )

0

I

2

4

6

8

NATURAL POROSITY, %

Fig. 3-13. Relation between carbonate content (and natural porosity) and crystal face index. (After Heaid et al., 1962, fig. 5, p. 295; courtesy J. Sed. Petrol.)

Zero indicates that the unblemished quartz faces were absent, whereas 10 is based on the highest proportion. Figure 3-12demonstrates that the unblemished crystal faces increased with porosity, indicating that the quartz faces in openings are well formed in contrast to the quartz terminations in contact with other grains. Heald et al. reasoned that if the observable quartz faces had been cleaned by leaching under natural conditions, then artificial leaching in the laboratory would have revealed more of the faces. The samples which were calcareous prior t o artificial leaching, however, had no higher crystal face index after leaching in comparison to the low-carbonate samples of equivalent porosity (Fig. 3-13). To rule out leaching before cementation, which is much more difficult to accomplish, Heald et al. considered the fact that most of the sandstone samples contain clastic carbonate grains (e.g., fossil fragments). If the porosity of the sandstone had been the result of leaching of the carbonate clasts, dissolution would have been widespread; but no evidence was found to support this. Although the clastic carbonate grains were typically mediumsand sized, open or cement-filled voids of this size were lacking, so that an origin of pores by removal of the clastic grains was not acceptable. Compaction could not have changed the shape and size of the voids because the more resistant detrital quaftz grains of the framework would have prevented it. In their section on porosity reduction, Heald et al. reasoned that because

DIAGENESIS OF SANDSTONES AND COMPACTION

107 SW *o. UC

It

N I,

-c

22

** I-

.?)

L 0 0

24

2 IT

28

t*

0

I1 . i

Fig. 3-14. Clastic constitutents, cement and porosity in the Oriskany Sandstone. (After Heald et al., 1962, fig. 7 , p. 297; courtesy J. Sed. Petrol.)

leaching was not a factor in the origin of pores, the porosity was controlled by incomplete cementation and degree of compaction, and presented evidence for their conclusion. The highly calcareous samples had a high primary porosity and consisted of well-sorted quartz grains and clastic carbonate constituents. Secondary cementation by sparite and quartz eliminated the porosity. The proportion of carbonate cement (= sparite) increased with primary carbonate clast content, most likely a result of solution and redeposition of the original carbonate. Variations in secondary quartz and carbonate sparite cements and pressure solution account for variation in the porosity of the more quartzose sandstone samples. In Fig. 3-14,the differences in the ratio (clastic grains)/ (cement + porosity) seem to be mainly due to the differences in the degree of chemical, pressure-solution compaction. This is supported by the number of contacts per grain and the degree of suturing, which are higher in specimens with higher proportions of clastic grains. The samples with 6% total carbonate content have relatively good porosities (Fig. 3-15).As Heald et al., however, pointed out, the relative importance of the two diagenetic factors in reducing porosity was difficult to assess. Heald and Renton (1966)reported on experiments carried out on sandstone cementation and its relation t o porosity changes. The tests were performed at 225"--360°C and at pressures from 2000 to 11,000 psi in hydrothermal reactors. The results, however, may apply to lower temperature and pressure conditions, if one considers geologic time as an additional influen-

K.H. WOLF AND G.V. CHILINGARIAN

108

6

P 0 R 0 S I T I (XI

I

Fig. 3-15. Relation between carbonate content and porosity in the Oriskany Sandstone, Kanawha County, West Virginia. (After Heald e t al., 1962, fig. 8, p. 298; courtesy J. Sed. Petrol.)

tial parameter. Particularly interesting were the observations on the control of cementation by variations in size, shape, angularity, and composition of the sand grains used during the tests, which demonstrated that the amount of artificial cementation was influenced by the original properties of the clastic components as well as by the variation in influx of cementing material. The initial rates of precipitation of the cement in fine-grained sand were greater than the rates in coarse-grained samples. As cementation proceeded, however, the coarser-grained sand cemented faster than the finer ones, because the permeabilities of the latter were appreciably reduced after a moderate degree of cementation. Heald and Renton’s experiments were divided into cementation of monocrystalline grains, polycrystalline grains, and mixtures of grains (e.g., quartz grains; lithic fragments of fine-grained sandstones, quartzites and cherts; and grains of arkose and micaceous quartzite).

109

DIAGENESIS OF SANDSTONES AND COMPACTION

Monocrystalline grains The following results were obtained by Heald and Renton from their experiments performed on monocrystalline grains: (1)The experiments on the relative growth rates of precipitated quartz on round quartz grains of different sizes are shown in Fig. 3-16.The gain in weight percentage due to cementation increased with decreasing grain size, but the ratio of the growth rates remained essentially constant. The greater rate of growth on the smaller grains is a reflection of the larger specific-surface area of the sand sample. (2) The rate of addition of secondarily precipitated cement is a function of the concentration of the solvent, but the ratio of the growth rates is independent of the concentration of the solution. (3)As Fig. 3-17 demonstrates, the relative growth rates of secondary quartz was the same in Na2C03 and NaOH solutions.

m

W

0

10

20

30

40

PERCENT GROWTH -STANDARDS

50

50

PERCENT GROWTH -STANDARDS

Fig. 3-16. Growth relationship between 9-10 mesh, 12-16 mesh, and 21-32 round quartz grains. (After Heald and Renton, 1966, fig. 2, p. 979.)

mesh

Fig. 3-17. Growth pf 12-16 mesh grains compared to standard grains in solutions of 0.03 M NaZC03 and 0.25 M NaOH. (After Heald and Renton, 1966, fig. 3, p. 980; courtesy J. sed. Petrol. )

K.H. WOLF AND G.V. GHILINGAREAN

110

)

-

P E R C E N T GROWTH STANDARDS

PERCENT GROWTH - STANDARDS

Fig. 3-18. Growth relationship between angular quartz grains of 9-10 mesh, 12-16 mesh and 21-32 mesh size. The 9-10 mesh grains and standard grains were contained in separate baskets, whereas in the charges of smaller grains, the standard grains were dispersed through the samples as internal sbandards. (After Heald and Renton, 1966, fig. 5, p. 981; courtesy J. Sed. Petrol.) Fig. 3-19. Growth relationship between aitgular and round grains of 12-16 (After Heald and Renton, 1966, fig. 6, p . 981; courtesy J. Sed. P e t r d . )

mesh size.

(4) It is noteworthy, that when 0.25-M solutions of Na2C03 were used, cryptocrystalline quartz was deposited, but in a reduced, 0.03-M solution normal quartz overgrowths were the result. (5) Figure 3-18 indicates that the precipitation of cement was faster on smaller angular than on larger angular grains. (6) Figures 3-19 and 3-20 prove that the relative amount of cement is consistently greater in the case of angular in contrast to round grains, when variations in circulation of the solutions and temperatures remain small.

Polycrystalline granular material

The conclusions to be drawn from Heald and Renton’s experiments on polycrystalline granular material are:

bIAGENESIS OF SANDSTONES AND COMPACTION

80

111

1

j-

PERCENT GROWTH - SEEDS

PERCENT GROWTH -BEREA POLYCRYSTALLINE

Fig. 3-20. Growth relationship between angular and round grains of 21-32 mesh size. (After Heald and Renton, 1966, fig. 7 , p. 981; courtesy J. Sed. Petrol.) Fig. 3-21. Growth relationship between granules of fine polycrystalline quartz and granules of coarse quartz from the Berea Sandstone. (After Heald and Renton, 1966, fig. 10, p. 983; courtesy J. Sed. Petrol.)

(1)The growth rate of quartz cement on very coarse grained quartz was nearly twice that of the growth rate on grains of fine-grained polycrystalline quartz fragments, as shown by the curve in Fig. 3-21 with all values along the y-axis being nearly twice that of the x-axis. (2) In experiments on monocrystalline quartz, quartzite and polycrystalline chert fragments of the same size and shape (and surface area, therefore), the growth rates in Fig. 3-22 demonstrate that (a) the rate for quartzite was low but increased with time. The slow rate of growth seems to be a reflection of the small crystal size of the polycrystalline grains. As the cement enlarged the small crystals with a cansequent increase in surface area, faster growth was promoted. (b) The lower values of growth rate for monocrystalline quartz was- the result of reduced circulation of the solution, because of porespace reduction by cementation. (c) The chei-ts grew even more slowly during the cementation experiments, the finer chert showing the lowest rate.

112

K.H. WOLF AND G.V. CHILINGARIAN

PERCENT GROWTH-POLYCRYSTALLINE QUARTZ

DURATION, Hours

Fig. 3-22. Growth relationship between monocrystalline and polycrystalline quartz grains.

1 = fine chert; 2 = medium chert; 3 = quartzite; 4 = monocrystalline quartz. (After Heald

and Renton, 1966, fig. 13, p. 985; courtesy J. Sed. Petrol.,)

Fig. 3-23. Porosity change during cementation of grains of quartzite and monocrystalline quartz. 1 = quartzite; 2 = quartz. (After Heald and Renton, 1966, fig. 14, p, 985; courtesy J. s e d . Petrol.)

(3)The porosity decreases during quartz-cement precipitation are presented in Figs. 3-23 and 3-24. Again, there are marked deviations between cementation of monocrystalline quartz and polycrystalline grains. Grains of hybrid composition

The results of the experiments by Heald and Renton on the precipitation of cement on grains of hybrid composition (i.e., mixtures) were: (1) The relative rates of cementation of an arkosic sand (= mixture of quartz and feldspar), in comparison with pure quartz sand, are illustrated in Fig. 3-25.Only 1-2%.of secondary quartz grew as very slender pesmatic crystals on the feldspar grains. Most of the quartz was precipitated on the

1 288

I U

432

OURATION, Hours

Fig. 3-24. Porosity change during cementation of grains of chert and monocrystalline quartz. 1 = chert; 2 = quartz. (After Heald and Renton, 1966, fig. 15, p. 985; courtesy J. Sed. Petrol.) W

I

I

I

I

1

I

a u)

20

40

-

60

80

PERCENT GROWTH PURE QUARTZ

Fig. 3-25. Growth relationship of arkose and micaceous quartzite compared to pure quartz. 1 = pure quhrtz; 2 = arkose; 3 = micaceous quartzite. (After Heald and Renton, 1966, fig. 16, p. 987; courtesy J. Sed. Petrol.)

K.H. WOLF AND G.V CHILINGARIAN

114

DURATION, Hours

DURATION, Hours

Fig. 3-26. Porosity change during cementation of arkose and pure quartz. 1 = quartz and feldspar (arkose); 2 = quartz. (After Heald and Renton, 1966, fig. r7,p . 987; courtesy J. Sed. P e t r o l . ) Fig. 3-27. Porosity change during cementation of micaceous quartzite and pure quartz. I = micaceous quartzite; 2 = quartz. (After Heald and Renton, 1966, fig. 18, p. 987; courtesy J. S e d . P e t r o l . )

quartz sand particles and was approximately proportional to the amount of quartz grains. (2) The changes in porosity during cementation of pure. quartz, arkosic sand, and micaceous quartzite sand, are given in Figs. 3-26 and 3-27. (3) The rates of cementation of sand composed of micaceous quartzite are shown in Fig. 3-25and the rate was much slower than in the case of quartzite, seemingly a result of the smaller surface area of quartz grains exposed. Geologic applications

Heald and Renton offered some geologic applications of th’eir experimental data:

DIAGENESIS OF SANDSTONES AND COMPACTION

115

(1)Well-sorted, coarse-grained sand will be cemented faster than finergrained sands as a consequence of greater influx of solutions into the more permeable coarser material. Where the invasion of solution is the same in both, however, the fine-grained sand is cemented more rapidly. To extend this conclusion to the theme of the present book on compaction, one must deduce from the above data that differential or preferential compaction of different sediments will depend on the numerous parameters, and may be prevented or modified by varying rates and degrees of cementation from one unit t o the next. (2) Heald and Renton showed that the medium- to coarse-grained sand 2 times faster than very coarse grained sand. Theoretically, very cemented : fine sand will cement 16 times faster than very coarse grained sand, as based on surface area differences. Consequently, if a bed or lens of coarse-grained sand is surrounded by a finer-grained sand, porosity reduction due to the precipitation of intergranular cement from solution would be quicker in the coarser-grained sediment. The rate of cementation would depend on the relative differences between the sand-grain populations. (3) As demonstrated from the data given by Heald and Renton, the rates

DURATION, Hours

Fig. 3-28. Porosity change during cementation of round and angular sand. 1 = round quartz; 2 = angular quartz. (After Heald and Renton, 1966, fig. 21, p. 989; courtesy J . Sed. Petrol.)

116

K.H. WOLF AND G.V. CHILINGARIAN

of the precipitation of cement vary with angularity of the sand particles: it proceeds considerably faster in angular than in round granular material of the same size grade. Figure 3-28illustrates that the higher porosity which is characteristic of angular sands is reduced to the same value of rounder sands after some chemical cementation takes place. As a result of differences in packing and because of the normally greater degree of initial compaction of angular sands, the combined effect of compaction and subsequent chemical cementation causes a more rapid decrease in porosity in angular sands. (4) Due to the different rates of chemical cementation of monocrystalline quartz sands, polycrystalline lithic sands, and polymineralogic (e.g., arkosic, micaceous, etc.) sands, there may have been different degrees of “lithification” from one bed t o another in a vertical sequence, as well as in the same bed horizontally, depending on the mineralogic composition of the original detrital material. These variations in susceptibility to cementation, among others, must be considered in studies of mechanical and chemical compaction on a regional scale. The studies of Levandowski and other investigators Levandowski et al. (1973)demonstrated the important role of cementation in petroleum geology, especially with regard to migration, accumulation, and storage of hydrocarbons in sandstones, because it controls porosity and permeability. Their work can serve as an example of detailed petrologic and geochemical principles being applied to regional diagenetic studies, which are discussed in a separate section below. A distinction was made between early, late and “differential” cementation by employing a paragenetic approach based on the age relationships of the cements, and the diagenetic features were then related to the paleohydrology of the basin in which the Permian Lyons Sandstone of Colorado accumulated. Similar techniques will have t o be used in ore genesis investigations in sedimentary and volcanic piles. The sandstone is quartz-rich with about 75% quartz and is bonded by a matrix and cement. As to the mineralogic composition, the individual beds vary from orthoquartzites to subarkoses. The composition of intergranular material is as follows: clay and chlorite = 1-1076; iron oxide = trace to 10%; secondary quartz = 0-2896; solid organic matter (i.e., petroleum residue) = up to 25%; anhydrite, which occurs as a common cement = 0-25%. Carbonate cement (calcite and dolomite) is also present and pyrite occurs in minor quantities as cubes and small patches. The paragenetic sequence of the above constituents is given in Fig. 3-29.Interesting to note is that solution and overgrowths on grains is absent where organic, iron oxide, and clay coatings are thickest and, apparently, have prevented the chemical interaction between the solutions and granular components. Anhydrite is

117

DIAGENESIS OF SANDSTONES AND COMPACTION EARLY

- -

PTLtDETRITUS GRAINS CLAY IRON OXIDE CEMENT CUARTZ

* LATE

~

SOLID ORGANIC MATTER

POSTDEPOSITIONAL STAGE

D

ANHYDRITE CARBONATE PYRiTE

--

-

Fig. 3-29. Paragenetic sequence of minerals in Lyons Sandstone, Denver Basin, Colorado. (After Levandowski et al., 1973, fig. 12, p. 2228; courtesy Am. Assoc. Pet. Geologists.)

associated with secondary quartz and dolomite, but rarely with calcite replacing quartz. The vertical and regional distributions of the quantitatively important cements are shown by cementation fucies in a cross-section and in a series of maps (e.g., Fig. 3-30). The cross-section gives part of the Lyons Sandstone in sw

NE

50110 ORGANIC MATTER

OM 'ORE

~

:Lx\z

CARBONATE A N D / M

ANMVWITE

PORE SPACE

0

MILES

Fig. 3-30. Distribution of cements and porosity in Lyons Sandstone. (After Levandowski et al., 1973, fig. 13, p. 2228; courtesy Am. Assoc. Pet. Geologists.)

118

K.H. WOLF AND G.V. CHILINGARIAN

which various cementing agents have been found (Le., quartz, organic matter, carbonate, and/or anhydrite) and the porosity was larger than 50%prior t o cementation (= pre-cement porosity). The extent to which Levandowski et al. have gone in their detailed petrologic evaluation is further demonstrated by their figs. 5, 7, 14 t o 18, 20, and 21, to which the reader is referred. They have prepared separate maps and/or cross-sections for: (1)tne distribution of red and gray iron oxide-bearing sandstones; (2) the distribution of quartz (see also Fig. 3-30 here); (3) the geographic variations of organic matter; and (4)the carbonate and anhydrite cements. In Fig. 3-30, it is shown that the distribution of quartz cement is independent of depth, but tends to increase towards the west. This build-up of quartz largely accounts for the permeability barrier. In the eastern part of the area, quartz grains are strongly etched, but little or no deposition of secondary quartz took place, indicating a removal or loss of S O 2 . The solid organic matter is present in a considerably larger area than oil, the former being closely associated with the gray Lyons sandstones, so that its regional distribution pattern is similar to that of these sandstones. The carbonate and anhydrite cements are shown to be concentrated near the top of the Lyons Sandstone, below the organic interval and below intervals containing a relatively large amount of quartz. Figure 3-30 gives their regional vertical distribution demonstrating the cement’s concentration at the top and base of the Lyons Sandstone and below the solid-organic matter zone, as well as at the flanks. The general build-up of the carbonate and anhydrite cements in the west, just east of the quartz concentration, is equally clear. Spectrochemical analysis of anhydrite demonstrated a regional Sr content variation in anhydrite (as a substitute of Ca*+).According t o Levandowski et al., the Sr content can be used empirically as criteria of anhydrite origin or type: (a) “Normal” or “primary” anhydrite present in bedded evaporite sequences tend to have a nearly constant Sr 102/CaS04ratio, because they formed in equilibrium with brines of similar composition. (b) The “secondarily redistributed” anhydrite (= “leached anhydrite”) has an Sr content that suggests frequent mobilization: Sr content of the normal anhydrite is lowered appreciably by extensive contact with formation waters. (c) “Secondarily introduced” anhydrite (= “enriched anhydrite ”) in veins and solution cavities is distinguished by higher than normal Sr content. The data by Levandowski et al. (1973) illustrate the enrichment of anhydrite in a northwest-southeast trending zone lying adjacent to and roughly parallel with the zone of high-Si02 cementation. It is associated with the known Lyons oil fields and the greatest enrichment occurs near these fields. The beds below the Lyons Formation and much of the Lyons sandstones to the east appear to contain leached anhydrite. The anhydrite of the cap rock

DIAGENESIS OF SANDSTONES AND COMPACTION

119

overlying the Lyons Sandstone is uniformly composed of normal anhydrite, suggesting that it was not affected by solution. Based on these regional distribution patterns, the formation waters seem to have moved upward into the Lyons Sandstone as well as laterally toward the edge of the basin because of the impermeable cap rock. This is supported by the petrographic study of the anhydrite itself and the fact that the greatest concentration of anhydrite occurs parallel with, and just basinward of, the permeability barrier (Fig. 3-30). As to the origin of the cements that are related to pressure and temperature in the lower ranges of values and, thus, being dependent on burial in the sedimentary basin, Levandowski et al. provided the following information. It should be noted at the beginning that the fresh-looking plagioclase with coatings of clay and chlorite in the sandstone samples suggested that burial metamorphism was absent or at a minimum and that the minerals making up the coatings were the result of diagenesis.

i WO

t

-

0

IO’IT~K

2

4

6

8

0

12

PH

Fig. 3-31. Solubility of quartz and amorphous silica as function of temperature. (From Siever, 1962, in Levandowski et al., 1973, fig. 25, p. 2238; courtesy Am. Assoc. Pet. Geologists.)

Fig. 3-32. Relation between pH and solubilities of calcite (after Correns, 1950), amorphous silica (after Krauskopf, 1958), and quartz (from data of Van Lier, 1959). Six areas are delineated: C = calcite; A = amorphous silica; Q = quartz; p = precipitates; d = dissolves; 1 = calcite in sea water at 2OoC; 2 = amorphous silica at approx. 25°C; 3 = quartz at 25OC. (Fiom Blatt, 1966, in: Levandowski et al., 1973, fig. 26, p. 2238; courtesy Am. Assoc. Pet. Geologists.)

K.H. WOLF AND G.V. CHILINGARIAN

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The source of the SiOz for the secondary quartz cement was fourfold: (1) solution of quartz sand grains; (2)deposition from sea water; (3)precipitation from connate water; (4) deposition from fluids derived from other beds. The fluids of types 3 and 4 sources could be wholly or partly made up of compaction fluids. As to the influence of overburden pressure, Kennedy (1950) showed that extremely high pressures, not usually encountered in sedimentary rock environments, are required to appreciably increase the solubility of quartz. On the other hand, Levandowski et al. (1973)found it reasonable to speculate that granulation and submicroscopic fracturing at grain contacts may produce some amorphous Si02, as does grinding of quartz in the laboratory (Siever, 1962). The interstitial water would preferentially dissolve the amorphous Si02 (Figs. 3-31and 3-32)and as the latter is removed in solution, the pressure along the grain contacts would produce new amorphous silica. Under natural conditions, other parameters in addition to the pressure would have to be considered, e.g., temperature and composition of pore fluids. Anhydrite and carbonate replacement of quartz releases Si02 to solutions that is then available for cementation. Figure 3-31 shows that the silica solubility increases with increasing temperature, whereas the CaCO solubility below 120°C and at constant Pcoz decreases with increasing tempera-

.20

c\

TEMPERATURE, O C

Fig. 3-33. Solubility of anhydrite in water, as function of temperature and pressure. 1 = pressure of 1000 bars; 2 = pressure of 500 bars; 3 = pressure of 1 0 0 bars; 4 = vapor pressure. (After Dickson et a t , 1963, in Levandowski et al., 1973, fig. 29, p. 2240; courtesy Am. Assoc. Pet. Geologists.)

DIAGENESIS OF SANDSTONES AND COMPACTION

121

ture. Anhydrite (= calcium sulfate) solubility also diminishes with increasing temperature (see discussion below and Fig. 3-33).If intrastratal fluids saturated with anhydrite and/or calcite are undergoing burial, temperature increases and CaCO and CaSO may precipitate, with concomittant dissolution of S O 2 . As Siever (1959)has mentioned, within the range of temperature t o which sedimentary rocks are subjected during subsidence of a basin, the temperature effect is much greater on the solubility of calcite and anhydrite than that of the pressure. The observation made by Levandowski et al. that S i 0 2 was moved from the basin center toward the margin where it was precipitated, indicates that the composition of the intrastratal fluids also influenced regional variations in cementation. As illustrated in Fig. 3-32 (e.g., Blatt, 1966), the calcitequartz replacement reactions are dependent on pH, e.g., dissolved SiOz tends to precipitate when the pH falls below 9.8. It may be proposed, therefore, that formation fluids in the sediments of the basin center were highly alkaline as a result of the presence of evaporites. The high alkalinity and the existing pressure and temperature conditions tended to increase the solubility of the quartz. Under late-diagenetic burial and compaction, the SiO2enriched waters moved horizontally underneath the tight anhydrite-rich unit. These basinal compaction fluids became mixed with less alkaline meteoric waters present along the basin margin. The resultant lowering in pH caused SiOz precipitation. As to the carbonate and sulfate cements, Levandowski et al. showed in their paragenetic sequence (Fig. 3-29)that they were the last to form, some evidence indicating that CaC03 may have been the latest of the two. The vertical and horizontal regional distribution of the last cement was at least partly controlled by the distribution of the earlier-formed intergranular pore fillings, e.g., quartz cement. The trace-element composition of the anhydrite suggested that the solutions moved from below and migrated laterally to the basin margins underneath the tight anhydrite cap rock. This took place during the compaction of the rock sequence. As shown in Fig. 3-33(Dickson et al., 1963), anhydrite solubility varies inversely with temperature and directly with pressure, whereas Fig. 3-34(Blount and Dickson, 1969)demonstrates that the anhydrite solubility is much greater in ionic NaCl solutions than in pure water at the same pressure and temperature. These investigators concluded that: (1) as a result of a combined temperature and pressure decrease that leads to a solubility increase, anhydrite will not precipitate from ascending saturated solutions; (2)at a constant temperature, anhydrite is precipitated from a saturated solution that migrated from a high-pressure to a low-pressure zone; and (3)anhydrite will precipitate from a high-ionic strength solutiov upon dilution by fresher water. Using the above data, Levandowski et al. reasoned that in the Lyons

K.H. WOLF AND G.V. CHILINGARIAN

122

to 0

250

1

1

500

750

I

1000

PRESSURE, Bars

Fig. 3-34.Anhydrite solubility as a function of pressure at several constant temperatures and NaCl concentrations. 1 = temperature = 100°C, NaCl concentration = 2M;2 = temperature = 100°C, NaCl concentration = 6 M ;3 = temperature = 15OoC,NaCl concentration = 6M; 4 = temperature = 2OO0C, NaCl concentration = 6M;5 = temperature = 2OO0C, NaCl concentration = 5M; 6 = temperature = 2OO0C, NaCl concentration = 4M;7 = temperature = 2OO0C, NaCl concentration = 2M; 8 = temperature = 100°C, in HzO;9 = temperature = 200°C, NaCl concentration = 1M;10 = temperature = 2OO0C, in HzO. (After Blount and Dickson, 1969,fig. 4,p. 235;courtesy Geochim. Cosmochim. A c t a . )

Sandstone sequence investigated by them, the formation fluids were highly alkaline, containing a large amount of sulfate. During compaction, the upward and, subsequently, laterally moving solutions moved towards the basin margins, where the pressure decrease resulted in precipitation of anhydrite and possibly some calcite. The theory of ionic impedance proposed by Fothergill (1955) was employed to explain the concentration of carbonate and anhydrite cement near a secondary quartz build-up that acted as a semi-permeable membrane behind which ions were concentrated on the high-pressure, basinward side. As compaction progressed, the formation fluids on the basinward side may have been subjected to higher pressures than on the landward side. Behind the semi-permeable membrane or barrier, ionic impedance progressively increased the concentration of Mg2+, Ca2+,and Sr2+ ions, until the solubility products of the respective minerals were reached and anhydrite and calcite were precipitated as open-space fillings. The sequence of cementation would have been: Sr-enriched anhydrite anhydrite -+ calcite. Figure 3-35 summarizes the geologic stages of the origin of the intergranular components present: --f

DIAGENESIS OF SANDSTONES AND COMPACTION R E PRESE N T A T I VE

GEOLOGIC

123

SETTING

STAGE 5 PRESENT- DAY

STAGE 4 ANHYDRITE AND CARBONATE

STAGE 3 PETROLEUM

STAGE I C L A Y AND IRON O X I D E

Fig. 3-35. Schematic diagrams illustrating cement development in Lyons Sandstone. (After Levandowski et al., 1973, fig. 31, p. 2242; courtesy Am. Assoc. Pet. Geologists.)

K.H. WOLF AND G.V. CHILINGARIAN

124

(1)Deposition of sands, with some detrital clay and iron oxide matrix. Basin*ard the Lyons sandstones are interbedded with limestones, dolostones, anhydrite (thick evaporite sequence) and shale. (2a) Accumulation of the anhydrite-rich unit (symbol: crosses) provided a relatively impermeable cap rock. (2b) Progressive accumulation of younger sediments increased the overburden, thus pressing out the connate water as compaction fluid. Migration was mainly lateral toward the basin margin within the sandstone underneath the anhydrite cap. The high alkalinity (pH) of the water caused SiOz dissolution and its subsequent reprecipitation as cement as pH was lowered. The silica solution predominated along the zone of the basinward sandstoneevaporite interbedding. On the other hand, near the western basin margin, the mixing of the alkaline with fresher water resulted in lowering of pH that caused SiO to precipitate. (3)Oil from the basinal shales and carbonate source rocks migrated into the Lyons Sandstone prior to complete compaction. The hydrocarbons moved laterally because of buoyant forces and the pressure gradient resulting from compaction. (4) With progressive burial, the overburden pressures led to the formation of solution effects, e.g., stylolites, with a consequent further expulsion of connate waters. These fluids dissolved CaS04 and CaC03 in older horizons and reprecipitated CaSO while migrating upwards and laterally. CARBONATE

4

A 0 0

LAUMONTITE

Jambwoo Sandstone Kioma Sandstone Westley Park Sondstonr

CHLORITE

LAUMONTITE

QUARTZ

CHLORITE

Fig. 3-36. Composition of the cement of sandstones from the Broughton Sandstone, Kiama, N.S.W.,Australia, based on carbonate, laumontite, chlorite, and quartz contents. Solid symbols represent micrometric analyses, whereas open symbols represent visual estimations; triangles = Jamberoo Sandstone; squares = Kiama Sandstone; circles = Westly Park Sandstone. (After Raam, 1968, fig. 6, p. 326; courtesy 3. Sed. Petrol.)

DIAGENESIS OF SANDSTONES AND COMPACTION

125

Raam (1968)described a terrigenous sequence in which the sandstones have a variety of matrices and cements of diagenetic origin (e.g., Fig. 3-36). Inasmuch as quartz and carbonates rarely are present together, two separate ternary diagrams were prepared. Zeolites (e.g., laumontite) and chlorite were frequently suggested to be of low-metamorphic or burial-metamorphic origin (e.g., Coombs, 1954, and Packham and Crook, 1960). Separate sections on burial metamorphism and related aspects are presented in this chapter. Raam attempted t o establish t o what degree the temperature and pressure increase was caused by overburden load, resulting in the formation of matrix and cement. He calculated that the unit was covered by sediments having a maximum thickness of only 2500-3000 ft, which is equivalent to a pressure of less than 200 atm and a temperature of about 35"C (assuming an average world geothermal gradient of loC/10O ft). Inasmuch as it was found elsewhere that laumontite forms at temperatures above 200" C and that although it can form in the presence of quartz at lower temperatures, i.e., 50--100"C, burial metamorphism could not have given rise to zeolites (e.g., laumontite), because they are present frequently in the absence of quartz. To find some additional support for his conclusion, he considered the rank of the coal in the rocks of the same basin and found that the maximum coal rank did not coincide with the general center of sediment accumulation, nor with the structural center of the basin. Raam, therefore, concluded that neither the authigenic minerals nor the high coal rank can be attributed entirely to burial metamorphism. He suggested that the physicochemical parameters of the early diagenetic environment played a more significant role. Hence, zeolitic reactions can occur at much lower temperatures and pressures than had previously been suggested. Hrabar and Potter (1969)gave an example of decementation. They found a sandstone body which was more permeable near the surface than its subsurface equivalents, because it had been decemented by ground water flowing downward through it into the underlying limestone (see Figs. 3-37A and 3-37B).The very high permeability values are due to the position of the sandstone body below the present erosion surface. Figure 3-38shows schematically the permeability distribution in an abandoned delta finger below an unconformity. The closer the sandstone to the unconformity, the greater its permeability. Important parameters controlling permeability were: (a) the length of time represented by the unconformity; (b) the intensity of decementation process; and (c) the extent of recementation after later burial. One might add another factor, i.e., any compaction that took place between cementation, decementation, and recementation could have reduced porosity and permeability. The formation of several generations of cement may be important in the .compaction history and should, therefore, be studied in detail.

K.H. WOLF AND G.V. CHILINGARIAN

126

Lithology and heah’ Mean Grain Size Bedding Facies (microns)

Horizontal Pernleability (100 millidarcies)

Percent Porosity

Fig. 3-37A. Vertical profile of bedding, grain size, permeability, and porosity, obtained from diamond drill cores. There is fining upward and vertical decline in permeability. Interval extends from Beech Creek Limestone to the top of Blue River Group. (After Hrabar and Potter, 1969,fig. 12;courtesy Am. Assoc. Pet, Geologists.)

Although the investigation of regional “diagenetic facies” is still in its infancy, a number of published examples clearly indicate a promising new field of petrology and sedimentary geochemistry. One variety of the diagenetic facies is based on the variation of different types of cements within sandstones (see also Wolf et al., 1967, pp. 128-130). Strakhov (1969, vol. 2,

OIO

20

50

too

zoo

500

low Moo

Jo

LOG PERMEABILITY, md

Fig. 3-37B. Stepwise cumulative curves of horizontal permeability values. There is a marked enhancement of permeability in outcrop. The 400 subsurface samples are from Bethel Sandstone in Midland field, Muhlenberg, Kentucky, whereas outcrop samples are from shallow drillhole in Owen County. There is an overall trend of permeability increase with depth. (After Hrabar and Potter, 1969,fig. 13, p. 2159;courtesy Am. Assoc. Pet. Geologists.)

DIAGENESIS OF SANDSTONES AND COMPACTION

2 .-

F

127

5600

c e .-

e 0

z-4000 0 t 3

I

I

I

\

\ - \

-

\

\

\

\

1

Fig. 3-39. Composition of the cementing minerals in various facies of Aalenian and Middle Jurassic deposits of Dagestan. Facies: A = alluvial; D = underwater deltaic; FB = foreset and bottomset beds of delta; M = marine; I = calcite; 2 = ankerite and siderite; 3 = quartz. (After G.N. Brovkov, 1958; in Strakhov, 1969, fig. 200, p. 513.)

pp. 513-515) mentioned the variations in distribution pattern of cement in sandstones and muddy sediments in different sedimentary depositional environmental facies (Fig. 3-39).He explained the variations to be the result of differences in fluid composition, differential migration of chemical elements during compaction, and related phenomena. Warner (1965, 1966) undertook a detailed study of the type, quantity, and distribution of cement, in addition to other parameters (Figs. 3-40,3-41, and 3-42),and demonstrated three stages of cementation by thin section analysis. He showed that the distribution pattern of the primary cement is related to the source rock and the distance of transportation from the source to the depositional milieu, whereas the secondary cements are more closely related to the structure of the region and the source rock, which suggests a ground water influence. Warner concluded that more investigations of this type are required to give information on sedimentary and tectonic structures, drainage direction of both surface and subsurface waters, permeability trends, distance of transportation influencing chemical differentiation, and source rock. The information gained would be of practical application not

1-1 1-

Outcrop limits of the Duchesne River Formation Contour representing percentage of cement Area of predominantly calcite cement Area of predominantly iron oxide. clayand silica cement

Fig. 3-40.Types of cements in the sandstones of the Duchesne River Formation. (After Warner, 1966,fig. 8, p. 952;courtesy Geol. SOC.Am.)

Contour representing percentage of grains showing secondary enlargement (silica cement) ...................... Synclinal axis in the Duchesne River Formation Fig. 3-41. Distribution of silica cement in Duchesne River Formation. (After Warner, 1966,fig. 9,p. 953;courtesy Geol. SOC.Am.)

DIAGENESIS OF SANDSTONES AND COMPACTION

-

129

Outcrop limits of the Duchesne River Formation

1-1

Contour representing the average number of trace elements

...

Sandstone faciesj 2.33;average

ratio

Mudstone facies,0.43:average

ratio

Fig. 3-42.Sandstone/mudstone ratio and facies of the Duchesne River Formation. (After Warner, 1966,fig. 13,p. 955;courtesy Geol. SOC.Am.)

Fig. 3-43. Carbonate mineral relationships in the CaC03-MgCOs-FeCO~ system. A. Composition of carbonate minerals in the Reedsville, Bald Eagle, and Juniata formations; two tie lines connecting mineral pairs are shown. B. Subsolidus relations at 400%. (After Rosenberg, 1960,1g67,based on laboratory study; in: Horowitz, 1971,fig. 1;courtesy J. Sed. Petrol.)

K.H. WOLF AND G.V. CHILINGARIAN

130

only in petroleum geology and hydrology, in general, but also in the understanding of the origin of copper and uranium "cements" in sandstones of the red-bed type (see Chapter 5 ) . The study of compactional history, together with textural and fabric investigations, may reveal the amount of fluids released, direction of fluid movement, and possible contribution of compaction fluids t o cement precipitation. Horowitz (1971) determined the stratigraphic variations in the ferrous iron content of the carbonates present in sandstones and used the data t o decipher the origin of these cements. Figure 3-43shows the general carbonsystem, whereas ate mineral relationships in the CaC0,-MgCO ,-FeCO Figs. 3-44 and 3-45illustrate the vertical changes in composition. Ankerite exhibited an upsection decline in ferrous iron content at four widely separated localities (Fig. 3-44),suggesting a vertical chemical gradient during its formation. Although Fez+ content in ankerite continued to increase downsection, total Fez+ in the carbonate system, i.e., ankerite plus calcite, actually decreased below the shoreline units (= Bald Eagle-Reedsville boundary),

,

REEDSVILLE

IOOO'

TYRONE

1' ' 1

8Od

I

600'

v,

LOYSBURG

1

0

t

I -

406

1

-

"

I

0

z

BEDFORD

P

n

u I

200'

u

100'

+ a

o

0,

a '3 a

Li

-100'

- 200'

NO ANK. ,

300

(Fe/Fe

I

30 0

+ Mg ) RATIO

Fig. 3-44. Stratigraphic variation in iron and magnesium content of ankerites at four localities in the Appalachian region of Pennsylvania. Color of rock units noted on side. Stratigraphic distance is shown in logarithmic intervals; zero mark shows the ReedsvilleBald Eagle boundary. (After Horowitz, 1971, fig. 2; courtesy J. Sed. Petrol.)

DIAGENESIS OF SANDSTONES AND COMPACTION

ca co,

131

Ca Mg (Cod.

Fig. 3-45. Approximate stratigraphic variation in bulk composition of carbonate cement. Letters relate positions in the stratigraphic column (left) to compositions on the ternary diagram. (After Horowitz, 1971, fig. 3; courtesy J. Sed. Petrol.)

because calcite became increasingly more abundant downsection (Fig. 3-45). Inasmuch as the decline of Fe2+ in carbonates above shoreline units corresponded to the successive upsection color change from dark grey to drab to red, it was concluded that the vertical chemical gradient reflected a change in the ogidation potential. Horowitz offered two hypotheses to explain his observations: (.1)successive facies differed in their Eh during sediment accumulation; or (2) diagenetic reducing solutions, originating from the lower marine units, reduced originally oxidized sediments with upsection effectiveness. Horowitn accepted the latter hypothesis as the more plausible one. The reducing fluids originated from the dark-colored marine Reedsville Shale and were expressed upward into continental red beds during compaction. Horor witz (1971) stated: “These solutions also may have carried the magnesium and calcium which combined with the reduced iron ,to form ankerite., since these two cations would. probably be more abundant in waters of marine origin. . . . That permeability variations influenced the migration paths of the reducing solutions can be inferred from the relationship between lithology and color. Proceeding westward from the depocenter, red beds are encountered first in the shalier (muddier) coastal plain facies (Fig. 3-46)and, more specifically, within shaly units of:this facies. Higher in the section, drab beds of the once more peimeable fluvial sandstane facies extend even further westward before passing into red beds. These observations suggested that reducing solutions favoured permeable sandstones. Displacement of most of the expressed water t o the west instead of eastward through the even ,more permeable conglomeratic facies wm atttibuted to the influx of meteoric water from the east and development of a hydrodynamic gradient to the west.” From the above discussions ‘one can conclude that the following factors must be considered in studying the cementation of sandstones by compaction fluids:

132 A

K.H. WOLF AND G.V. CHILINGARIAN A'

UNCONFORMITY

Fig. 3-46. Model proposed to show the migration paths of reducing fluids and explain the origin of the red-drab color boundary between the Bald Eagle and Juniata formations. 1 = reds beds, Juniata Formation; 2 = black shale, Reedsville Formation; 3 = drab beds, Bald Eagle Formation; 4 = direction of water movement. (After Horowitz, 1971, fig. 4; courtesy J. Sed. Petrol. )

(1)The distribution of sedimentary facies (i.e., conglomerates, sandstones, clayey deposits, limestones, dolomites, and evaporites), which deterniines the regional variations in primary porosity, original amounts of fluids trapped within the sediments and, therefore, the potential rate and degree of compaction to be expected on a theoretical basis. (2) The mineralogic composition of sedimentary particles, which determines the rate of release and composition of the fluids from clay minerals, for example, to form compaction solutions (see separate sections in this chapter, pp. 290-303). (3) The textures and labrics, which indicate the mechanisms, degree, and regional variation of compaction. (4) The diagenetic effects (e.g., genesis of cements and matrices; removal of heavy minerals by chemical dissolution; paragenetic relationships; alterations of clays). (5) The structural history of the sedimentary basin, which controls the amount and rate of subsidence and, therefore, the rate of sediment accumulation and the rate of compaction. The latter, in turn, controls the rate and direction of compaction fluid movements. The regional structure will also control later ground water. migration, which may cause decementation and recementation that could erase or obliterate the earlier-f ormed diagenetic features. To the knowledge of the writers, a total and all-inclusive regional

DIAGENESIS OF SANDSTONES AND COMPACTION

133

petrologic and geochemical investigation of the compaction history of a sedimentary basin has not been undertaken as yet. Most investigators, who have considered compaction in their stratigraphic and environmental examinations of sedimentary piles, have employed one or two particular techniques, such as studying textural changes in a vertical profile and/or variations in cementation, that resulted in data satisfactory for their particular needs and goals. Inasmuch as numerous different techniques and concepts have been developed and are available for compaction studies, they should be combined in the future in a deliberately comprehensive examination of the three-dimensional compaction history of a sedimentary basin. Such investigations could be particularly attractive because of their practical applicability in the petroliferous basins and sedimentary ore districts, and may also assist in unravelling the possible relationships between hydrocarbons and ores in the same basin (see Chapter 5). TEXTURES* RESULTING FROM COMPACTION

Geologists require details on the textures and their paragenetic relationships, in addition to data on composition, to be able to unfold the geologic history of sedimentary rocks. Much data are now available on the textures of the clastic rocks, especially the sandstones. The textures have their start with primary depositional characteristics and grade without definite boundaries into the secondary features, among which are those formed by compaction as a result of overburden pressures. The strictly primary and secondary-diagenetic textures, in turn, grade into catagenetic and metamorphic textures, fabrics, and structures. Thus, it is essential to be familiar with all types of petrographic characteristics in sedimentary petrology. There is hardly any concept in the interpretation of the origin of sediments that does not include their textural properties, so that for a proper comprehension of the various sections of this chapter it is necessary to have some understanding of the textures of sandstones. To name but one example, Modarresi and Griffiths (1963)have demonstrated by the use of statistical petrographic analyses of reservoir rocks that the textures of the framework of sandstones controlled subsequently developed silica and carbonate distribution. Although textures and fabrics are fundamental units of the larger-scale structures in sedimentary deposits, the latter are not considered here. Large features resulting from compaction and which, therefore, fall into the group of secondary structures, are also bypassed in this section, but a list with *The term “texture”’as used here includes those features that have been called “fabrics” by others, but excludes the large-scale structures.

K.H. WOLF AND G.V. CHILINGARIAN

134

pertaining references is provided at the end of this chapter. Most of the information obtained on textures and fabrics comes, from studies utilizing the petrographic microscope, but the use of the electron microscope and X-ray textuameter is rapidly increasing. Sippel (1968) employed luminescence petrography in the study of sandstone diagenesis and was able to: (a) distinguish detrital quartz from quartz overgrowths; (b) recognize fracture-healed quartz grains as well as features resulting from crushing; and (c) distinguish recrystallization from polycrystalline features inherited from themouroe area. All these textures are frequently not recognieable under a normally-equipped petrographic microscope. Pettijohn\et 4.(1972) presented a summary of the fabrics of sands and sandstones. They stated that the arrangement of grains to form an aggregate reflects: (a) manner of deposition at the time of sediment accumulation; (b) grain size; (c) sorting; (d) shape of the clasts; and (e) physical. and chemical compaction. The measurements of these parameters, together with other information, such as mineralogic composition, type and abundance of fossils, structures, and regional relationships of the various lithologic units, enables

SbTURED GRAINS

CONCAW-CONVEX CONTACT

POINT CONTACXLONG CONTACT

FLdATlMG GRAINS

LlNE .OF TRAVERSE

Fig. 3-47. Definition sketch of fabric terminology: quartz (white), mica (lined), and matrix (stippled). Illustration of the application of the quantitative fabric indices listed in Table 3-XVIIL. (After -Pe&ijohn et al., 1972, fig. .3-10;p. 91; courtesy Springer, New York.)

D M E N E S I S OF SANDSTONES AND COMPACTION

135

one to infer the primary and secondary rock formation regimes. The above information can also serve as a guide in estimating the crushing and bearing strength of sediments and rocks, as illustrated in some of the chapters of both Volumes I and I1 of this book. Pettijohn et al. (1972) offered a summary of the qualitative and quantitative teFms and indices that are used in describing the grain-to-grain relations. They have pointed out that the terminology and methodology are available now for detailed textural investigations, but there is a shortage of systematic mapping studies that could serve as models. Table 3-XVIII presents the TABLE 3-XVIII Terms used to specify sandstone fabric (after Pettijohn et al., 1972, table 3-4, p. 90) Qualitative Concauo-convex contact (Taylor, 1950, p. 707): one that appears as a curved line in the plane of section Fixed margin (Allen, 1962, p. 678): that part of a grain in contact with another in the plane of section Fixed grain (Allen, 1962, p. 678): fixed margin exceeds free margin Floating grain: no contacts with other grains in the plane of section Framework fraction: the stress-transmitting portion of a sand Free margin (Allen, 1962, p. 678): that part of grain not in contact with other grains in the plane of section Free grain (Allen, 1962, p. 678): free hargin exceeds fixed margin Long contact (Taylor, 1950, p. 707): a contact that appears as a straight line in the plane of section Packing (Kahn, 1956, p. 390): mutual spatial relationships among grains Sutured contact: mutual stylolitic interpenetration' of two or more grains Tangential contact (Taylor, 1950, p. 707): one that appears as a point in the plane of section Quantitative Condensation index (Allen, 1962, p. 678): ratio of percentage of fixed rock fragments t o percentage of free grains Contact index : number of contacts per grain Horizontal packing intercept (Mellon, 1964, fig. 7): average horizontal distance between framework grains Packing density (Kahn, 1956, p. 390): length of grains intercepted divided by length of traverse X 100 Packihg index (Emery and Griffiths, 1954, p. 71): the product of the number of quartz to quartz contacts per traverse and the average quartz diameter, the product being divided by the total length of traverse Packing proximity (Kahn, 1956, p. 390): number of grain-to-grain contacts divided by total number of contacts of all kinds (grain-to-matrix and grain-to-cement)X 100. Vertical packing intercept (Mellon, 1964, fig. 7): average vertical distance between framework grains

136

K.H. WOLF AND G.V. CHILINGARIAN

TABLE 3-XIX Sandstone fabric (modified from Adams, 1964, table 1, after Pettijohn e t al., 1972, table 3-5, p. 92) A. Grain-to-grain relations specified chiefly by qualitative observation supplemented by some quantitative indices on non-oriented samples; chiefly used to interpret and predict reservoir porosity ( 1 ) Much pressure solution: many sutured contacts, grain-to-grain contacts large, packing density high, and no porosity and cement (21 Moderate pressure solution : some sutured contacts, principally equidimensional grains, chiefly with concave-convex contacts; cement is mostly quartz overgrowths with or without minor amounts of carbonate or clay; little porosity (3) Minor pressure solution: mostly original grain outlines with long and tangential contacts; low t o moderate number of grain-to-grain contacts and moderate packing density; may be either well or poorly cemented by quartz overgrowths, clay, and carbonate; may have moderate porosity, if poorly cemented ( 4 ) No pressure solution: chiefly original grain outlines with either tangential contacts or floating grains. Number of grain-to-grain contacts is low as is packing density; cements are mostly carbonate or clay; porosity is high when cementation is limited B. Orientation of framework specified by quantitative measurement of oriented samples; porosity and cementation are independent of orientation; chiefly used t o determine current direction in undeformed sediments ( 1 ) Particulate methods: visual, direct measurement of either long axes or apparent long axes of framework grains, usually in thin section (2) Aggregate methods: measurement of a bulk geophysical property by an appropriate black box that can be correlated with the orientation,of the framework grains

qualitative and quantitative terms employed in textural investigations of sandstones, whereas Fig. 3-47 illustrates the quantitative terminology for textures. One should note that Griffiths (1967b) used a somewhat different nomenclature as given in Table 3-XX.’ One example, based on actual observations showing the relationships among fabric, pressure solution, and cementing agents, is presented in Table 3-XIX. Griffiths (1967a) and Blatt (1966) stated that packing, a parameter of particular importance in compaction, depends on: (1)mode of sediment deposition; (2) orientation of the particles, (3) grain size; (4) grain shape; (5) range in grain size (= sorting); (6) amount of clay matrix; (7) mineralogic composition; ( 8 ) overburden pressure and degree of compaction; and (9) time of precipitation of cementing agents. Both cementation and compaction convert loose sediments into consolidated rocks, and are part of what is usually called “lithification”. The amount of compaction of sand, according to Blatt (1966) depends on: (a) mean grain size; (b) sorting; (c) amount of argillaceous material present; and (d) other factors that are, however, believed to be of minor importance. In

DIAGENESIS OF SANDSTONES AND COMPACTION

137

TABLE 3-XX Classification of the configuration of contacts (after Griffiths, 196713, table 8.6, p. 173) Kind of contact

Symbol

Diagrammatic appearance

Class No.

~~~

Floating

F

0

Tangent

T

1

Long

L

2

Complete

C

3

Sutured or serrated

S

4

particular, the influence of a fine-grained matrix on compaction requires special attention in future research. The problem is accentuated by the fact that there are two types of matrices, namely, a detrital, primary variety and a diagenetic-metamorphic, secondary type. Determinations of the relationships among relative proportions and types of the grains, matrix and cement and their ratios, on one hand, and the style, amount and rate of compaction, on the other, necessitate a precise terminology in the study of the matrices. Dickinson's (1970) publication is a very significant one from this point of

K.H. WOLF AND G.V. CHILINGARIAN

138

.,....

. ......... . . . I.

...

,.

..........

.

\ . . . .. .\ . . . . . . .

.....

b

a

0.

.: . .

C

..'

. .. .

... ...

. .. . . . . . . . . / . . .. . .. . :

d

.

.

f

e

;:,*.;

......... ..

. . . . .: ..: .

,./_.

.. : . ..... . .:

-.

:,..... . '!>. .I_.

..

5

5.

. .

..

..

............... . ..............

h

I

k

I

...

.; . .:,..:.

..

%.

Fig. 3-48. Grain-enlargement, pressure-solution, and micro-drusy (open-space filling) textures, (After Glover, 1963, fig. 5, p.5l;courtesyJ. R . SOC.W. Aust.) 1. Enlargement textures. Simple enlargement textures: ( a ) Calcarenite showing sparry calcite in crywtallographic continuity with crinoid debris. Stippled fraghents are calcare-

DIAGENESIS OF SANDSTONES AND COMPACTION

139

view as it summarizes his own as well as the ideas of other investigators. He offered useful terms for matrices of four different origins: proto-, ortho-, epi-, and pseudo-matrix. (For a related terminology on carbonate sediments, see Wolf and Conolly, 1965.) The papers by Dapples (1962,1971, 1972) on the diagenesis of sandstones should also be consulted. As pointed out already, in the investigation of compaction, the petrologist must be familiar with all types of genetic textural varieties, i.e., both of non-compaction and compaction origin. For this reason, summaries on textures based on the studies by Glover (1963) and Strakhov (1957), are given here in the form of diagrams (Figs. 3-48, 3-49, 3-50 and 3-51). The former investigator offered five basic diagenetic textural groups of general application, some of which can be directly attributed to compaction, whereas others could be either indirectly the result of the interaction with compaction fluids or could be due to other phenomena. Strakhov’s diagrams are particularly useful in the investigation of cements and matrices in conjunction with the information supplied by Dapples (1962, 1971, 1972) and Dickinson (1970). Much information of interest t o sedimentary petrologists can be obtained

ous pellets. ( b ) Quartz sandstone with quartz outgrowths crystallographically continuous with clastic cores. Note faces. Lightly stippled areas represent pores. (c) Quartz sandstone with pores completely occupied with secondary quartz. Some outgrowths bounded by plane surfaces, some not. No sutured boundaries. Indentation textures: ( d )Dolomitic sandy marl in which dolomite is partly surrounded by, or has partly penetrated, quartz outgrowths. Texture does not reveal whether quartz or dolomite grew first. ( e ) Dolomitic sandy marl, same as ( d ) , except one of the dolomite grains is moulded onto a clastic quartz core. Dolomite, therefore, preceded secondary quartz. Enclosure texture: ( f ) Dolomitic sandstone with dolomite completely enclosed by quartz. Dolomite, therefore, formed first, and order is confirmed by moulded dolomite (upper center). Note how a moulded dolomite has retreated marginally (left center) due to slight solution during silicification. 2. Pressure-solution textures. ( g ) Quartz sandstone with clastic grains showing sutured boundaries due to compaction, deformation or both. Secondary quartz(s) which fills voids, may have come partly or completely from quartz dissolved along sutured contacts. (h) Calcarenite with microstylolite due to compaction. Microstylolite outlined by ironstained argillaceous matter and small quartz grains, both insoluble in the particular conditions of its formation here. Sparse distribution of quartz in rock suggests compaction equivalent to field of view. (i) Quartzite with sutured boundaries between outgrowths due to deformation after diagenesis. As much a metamorphic as a diagenetic texture. 3. Micro-drusy textures (or pore-filling textures). Simple micro-drusy textures: (j)Calcarenite partly cemented with fibrous calcite. Fibers are elongated normal to grain boundaries. (k) Calcarenite completely cemented with sparry calcite. Long axes of calcite crystals are normal to grain boundaries. ( 1 ) Lithic (volcanic) sandstone cemented by fibrous chlorite. Texture basically the same as in 0’) and (k).

K.H. WOLF AND G.V. CHILINGARIAN

140

a

C

f

b

d

e

n

Fig. 3-49. Composite micro-drusy, reorganization, and replacement textures. (After Glover, 1963, fig. 6, p. 53; cqurtesy J. R. SOC. W.A u s t . ) Composite micro-drusy textures: ( a ) Lithic (volcanic) sandstone with pores filled by three minerals which are, from the outside, a micaceous mineral, chlorite (stippled), and feldspar. Note how minute chlorite

DIAGENESIS OF SANDSTONES AND COMPACTION

141

from the literature on soils (e.g., Brewer, 1964), but only a few serious attempts have been made in this direction. The publications by civil engineers on compaction and compressibility of soils have also been only occasionally consulted by sedimentologists. Pedologists, for example, use textural terminologies and genetic concepts that may be applicable not only to paleosoil investigations (e.g., Yaalon, 1971),but also in sedimentological studies. The results of the work on artificial compaction of soils may throw some light on the compaction of sediments. The textural data supplied by soil specialists will eventually allow one to establish criteria enabling a petrologist to differentiate genuine sediments from paleosoil per se as well as permit a formulation of a list of textural differences and similarities. Soils are often composed of a whole gamut of individual grains of sand, silt, and flocculated clay minerals arranged in an arching skeleton (Fig. 3-52) enclosing large voids, termed “honey-comb” fabric (Casagrande, 1940, p. 85; see also Gillott, 1968, fig. 27). With the increasing use of electron-microscope techniques in determining the origin of the different types of matrices of sandstones, it is suggested that the petrologist may find the pedological literature a fruitful mine for ideas in attempts to differentiate between the clay matrices. Figure 3-53,for example, shows the various modes of textural associations of clay minerals as used by pedologists. Similar modes should be serrations are directed inward. The micaceous mineral may be a reconstituted clay film on the clastic grains. The order of formation was: (1)micaceous mineral (or its precursor), (2) chlorite, and (3) feldspar. (Based on a diagenetic sequence in Arrowsmith Sandstone.) (b) Lithic (volcanic) sandstone showing the following diagenetic sequence: (1)micaceous mineral; (2) quartz; (3) chlorite; (4) quartz (note euhedrism); (5) feldspar. (Based on a diagenetic sequence in Arrowsmith Sandstone.) Reorganization textures: ( c ) Claystone with vermicular kaolinite crystals. Fragility of the crystals is a proof of in situ formation. ( d ) Graywacke with chlorite-sericite matrix, formed from clay-sized detritus. The new flaky minerals penetrate margins of clastic fragments, making this partly a replacement texture also. (e) Fontainebleau Sandstone in which calcite has reorganized to form large crystals. Replacement textures: ( f )Calcilutite partly replaced by quartz euhedra with calcareous inclusions. The long fragment is calcite. ( g ) Shelly limestone with shells partly replaced by chalcedony; matrix is dolomitized. The dolomite has zonal inclusions, see 0’). Unreplaced shelly material has recrystallized. (h) Quartz sandstone cemented by sparry barite. The original pyritic and argillaceous matrix is represented by patches of argillaceous impurity rind isolated pyrite grains. (i) Ferruginous quartz sandstone in which quartz grains are apparently corroded by the ferruginous matrix. One quartz grain shows an outgrowth product of an earlier diagenetic phase. The texture resembles that where carbonate corrodes quartz; many such sandstones may originate as a result of replacement of calcite by iron oxide. 0’) A complex but fairly common texture, in which dolomite has partly replaced the matrix of a sandy marl. Carbonaceous and argillaceous inclusions form a dark zone in each dblomite rhomb. There has been later recrystallization of the matrix to sparry calcite, with expulsion of impurities.

K.H. WOLF AND G.V. CHILINGARIAN

142

. . .;

.

,.

Fig. 3-50. Authigenic enlargement of quartz in sandstone from the Birdrong Formation (Rough Range Bore No. 1, core 7 , 3,633-3,636 ft). The stippled areas are partly filled with clay-sized material. Quartz crystal faces are indicated by p (prism) and r (rhombohedron). The sketch is slightly idealized to show faces clearly. Width o f field = 0.6 mm. (After Glover, 1963, fig. 2, p. 39; courtesy J. R. SOC.W. Aust.)

expected in the matrices of sandstones and in mudstones, in particular if the clay is of diagenetic origin and in an uncompacted state. To better understand the very early compaction mechanisms of sediments (and physical and chemical diagenesis, in general), the textures of detrital, flocculated and chemically precipitated clay accumulations must be thoroughly studied using electron microscopy. Griffiths (1961) pointed out that some of the most important aspects of sedimentary petrography are definition and evaluation of procedure and an outline of a uniform methodology. He also stated that relatively few detailed analyses of sedimentary rocks exist and most of those that do exist are qualitative rather than quantitative. Griffiths offered a “conceptual definition” of a sedimentary rock by specifying five properties, which are to be considered also in detailed investigations related to physical and chemical compaction: (1)proportions of different kinds of grains (= m ) ; (2) sizes of grains (= s); (3) shapes of grains (= s h ) ; (4)orientation of grains (= 0);and (5) mutual arrangement or packing of grains ( = p ) . In several publications and books, Griffiths then has offered the following formula that depicts the properties of a sediment: P = f(m,s,sh,o,p), where P = unique index and f = function. Griffiths (1961, 1967a) discussed the complex interrelationships among these properties and described procedures for specification of each property, which when expressed as numbers, can be substituted in the above formula. The resulting value of P is a unique description of the rock specimen. In a multivariate system, the properties may be interdependent, dependent, or independent, as Shown by the three-variable system in Fig. 3-54 (A& and C ) : ( A ) size changes with shape, showing a direct linear depen-

143

DIAGENESIS OF SANDSTONES AND COMPACTION f

4

5

7

6 d

6 C

d

a

Y

€

4

71

Fig. 3-51. Textural types of cements. 1-8: types of cement association with detrital grains; a-e: argillaceous cements; and a-: cements of chemical origin. (After Strakhov, 1957, fig. 20.)

144

K.H. WOLF AND G.V. CHILINGARIAN

Fig. 3-52.Proposed fabric of clay soils. Honeycomb-structure flocculated soil. (After Casagrande, 1940,p. 85;in: Gillott, 1968,fig. 27;courtesy Elsevier, Amsterdam.)

Fig. 3-53.Modes of particle association in clay suspensions, and terminology. A. “Dispersed” and “deflocculated”; B. “aggregated” but “deflocculated” (face-to-face association, or parallel or oriented aggregation); C. edge-to-face flocculated but “dispersed”; D. edge-to-edge flocculated but “dispersed”; E. edge-to-face flocculated and “aggregated”; F. edge-to-edge” flocculated and “aggregated”; G. edge-to-face and edge-tosdge flocculated and “aggregated”. (After Van Olphen, 1963, p. 94; in Gillott, 1968, fig. 37;courtesy Elsevier, Amsterdam.)

*-I/

145

DIAGENESIS OF SANDSTONES AND COMPACTION A) DEPENDENCE

A

Y = a + p x

W

U

a

a

I

0

a

0

= comt k ; 8'1

0

0 0

SIZE,X Size changes with shape; lineor dependence

3 INTERDEPENDENCE

8 ) INDEPENDENCE

I

0

SI2E.X Size

0

0

0

independent of shape

t

to shope but relationship with size changes

SiZE,X Size reloted

Fig. 3-54. Interrelationship among the properties of sediments. (After Griffiths, 1961, fig. 1, by permission of The University of Chicago Press, copyright @ 1961 by the University of Chicago.)

Fig. 3-65.Geometrical illustration of three-factor experiment and its interactions. Main effect = solid lines; first-order interactions = broken lines; and second-order interaction = dotted line. (After'Griffiths, 1961,p. 493, by permission of The University of Chicago Press; copyright 0 1961 by the University of Chicago.)

146

K.H. WOLF AND G.V. CHILINGARIAN

TABLE 3-XXI Three-factor experimental design t o illustrate sources of variation (after Griffiths, 1961, table 1, p. 4 9 1 ) Source of variation: Size Shape Orientation

three main effects

Sizeahape Size-orientation Shape-orientation

three first-order interactions

Size-shape-orientation

one second-order interaction

dence; ( B ) size is independent of shape; and (C) size is related to shape, but the sizeshape relationship is not the same in all size ranges. Griffiths (1961) then proceeded to a three-variable system which is even more complex; the sources of variation are summarized in Table 3-XXI and graphically represented in Fig. 3-55. In a discussion of the application of his equation, Griffiths mentioned that he has used it successfully in: (a) detecting a favorable change from barren sandstones to potential oil reservoir rocks, and (b) differentiating ore-bearing from barren sandstones in the uranium-containing sandstones of the Colorado Plateau-type of mineralization. His approach should also find application in detailed work on compaction where precise description of all variables is fundamental.. Using “bedding” as an example, the problem of defining the microstructure of a sedimentary rock can be demonstrated. According to Griffiths, bedding represents heterogeneity or non-randomness of arrangement of the constituent elements. Bedding may be due to: (1) change of composition (Fig. 3-56,]); (2) to grain-size variation (Fig. 3-56,2) if composition is constant; (3) to grain-shape variation (Fig. 3-56,3) if both composition and grain size are constant; (4)to a change in orientation (imbrication) (Fig. 3-56,4), if composition, grain size and grain shape are maintained constant; and (5) to differential packing (Fig. 3-56,5), if all four properties remain constant. The question arises here as to how changes in one or all of these five parameters would affect chemical and physical compaction. Most studies made so far in this area have been rather simple and were based on one or two variables; however, in more recent studies, attempts are being made to consider more variables and their interdependencies. If subtle differences in the above-mentioned grain parameters influence the mode, degree, and rate of compaction of sediments, and if a set of more or less constant characteris-

DIAGENESIS OF SANDSTONES AND COMPACTION ~

G OF C OEM P O S ~ T ~ O N

I3.CHANGE

OF SHAPE

147

1 2 . CHANGE OF SIZE

I 4 . C H A N C E OF ORIENTATION

I

Fig. 3-56. The fundamental basis of “bedded” microstructure. (After Griffiths, 1961, fig. 3, p. 495; by permission of The University of Chicago Press, copyright @ 1961 by the University of Chicago.)

tics that vary between certain limits can be established for each sandstone group, then each type of sandstone can be characterized by certain differences in physical and chemical diagenesis, including compaction. Although a relatively abundant data is available already on the differences and similarities between arkoses, graywackes and quartzites (some of the differences are inherited from the definitions of these three petrographic terms and the arbitrary boundaries set by different researchers), many aspects remain to be investigated. One example is that described by Griffiths (1967b) on the mean and variance of roundness in the three basic sandstone types (Fig. 3-57): the arkoses show a wide variability in variance for a very small range in mean roundness, whereas the graywackes and quartzites exhibit a trend in which the variance increases with increase in mean roundness. The scatter of the data for the quartzites is largest, i.e., they differ in both sphericity and roundness considerably (see the very inclined line in Fig. 3-57). These differences are not the result of differences of the various rock types in the source area, but reflect the degree of selective shape-sorting of the grains: the arkoses are all near-source sediments, the graywackes vary in both texture and composition, and the quartzites are of several different textural types. The near-source sediments were subject t o relatively little selective sorting and, therefore, show wide variability within samples but little among sample means. On the other hand, well-transported sediments tend to exhibit wide variability in sample means and less variability within samples. One example

148

K.H. WOLF AND G.V. CHILINGARIAN

/-

---

TRENDS INSERTED BV EYE

@,a,@,ROCK-TYPE

MEANS

Fig. 3-57. Mean and variance of roundness measurements. (From Curray, 1949; in: Griffiths, 1967b, p. 139; by permission of McGraw-Hill, New York, 0 1967 McGraw-Hill, New York.)

of the change of sphericity and roundness with distance of transportation of the grains from the source area is presented in Fig. 3-58.The above geological occurrences are to be considered “ideal” situations and should be used as “models” in the search of a number of possible exceptions. In much of the theoretical considerations of the porosity, permeability, and depositional fabrics and textures of sandstones, most investigators use spherical particles as the fundamental, ideal constituents of the sediments. Inasmuch as compaction features and the behavior of sand grains under stress, for example, are directly related to the initial characteristics of the

F

i DISTANCE, Miles

Fig. 3-58. Form of the function representing the relationship of change in particle shape with distance (or time). 1 = analogous to .Spey River (after Mackie, 1897); 2 = analogous to Mississippi River (after Russell and Taylor, 1937). (Based on Krumbein, 1941; in: Griffiths, 1967b, fig. 6-16; by permission of McGraw-Hill, New York, copyright 0 1967 McGraw-Hill, New York.)

DIAGENESIS OF SANDSTONES AND COMPACTION

149

grains, the theoretical evaluations of the secondary modifications are also based on the assumption that constituent particles are ideal spheres. Allen (1969, 1970), however, has pointed out in several papers that theory and experiments based on the sphere as the ideal particle, fail to account satisfactorily for the concentration G f solids observed in unconsolidated natural sands. Allen (1969, p. 309)stated that “when the prolate spheroid of moderate axial ratio is substituted for the sphere, a theory which is satisfied by observation becomes possible, for the reason that equal spheroids can be regularly packed in both close and open ways”. Allen pointed out that the freshly deposited natural sands have a relatively low particle concentration, even though theory based on the sphere as the ideal sedimentary particle predicts for them a relatively high concentration of solids. Natural sands are chiefly deposited either as laminae on horizontal or nearly horizontal surfaces or as steeply-inclined laminae on slip faces or ripples and/or dunes (see results of numerous experiments by Jopling, 1963, 1965a,b, 1966, 1967, for example). The concentration of grains, C , which is equal to (space occupied by solids)/(total space), ranges between 0.50 and 0.65 in the above two cases. The variable C depends on the average grain size, degree of sorting, medium of deposition and intensity of deposition, and the porosity is equal to (1-C). According to Allen (1969, pp. 309-310), the range of possible particle concentration C, obtained experimentally by using haphazardly packed equal solid spheres is roughly between 0.60 and 0.64. Theory based on the sphere as the ideal sedimentary particle, as mentioned in most text and reference books, predicts C to increase with increasing range of sizes of different sphere populations present. Consequently, C of natural sands should be distinctly larger than the possible range of C for equal spheres, because natural sands are composed of grains having an appreciable range of sizes. Observations do not agree with this, because the lower limit of C for natural sands is about 0.5, which is below the lower possible limit of C for haphazardly

mm Cubic Packing

Orthorhombic PockinQ

PACKING Ill

Tetragonol Packing

Rhombohedra1 Pocking

Fig. 3-59. The four dose packings of equal prolate spheroids when the major axes lie in the planes of the layers. (After Allen, 1969, fig. 1, p. 312; courtesy Geol. Mag.)

K.H. WOLF AND G.V. CHILINGARIAN

150

packed equal spheres by 0.1. The upper limit of C for natural sands deviates little from the upper limit for haphazardly accumulated spheres. Allen attributed this to the use of the sphere as an ideal sedimentary particle. As shown by numerous investigations, the natural sand particle is approximated by a triaxial ellipsoid with a long axis about 1.5 times longer than the intermediate axis and about 2.0 times longer than the short axis. Allen (1969, p. 310) continued to explain that equal solid ellipsoids can be regularly packed in more different ways than equal solid spheres, because these particles can be differently oriented in space, which need not be the same for all the ellipsoids (Figs. 3-59, 3-60, 3-61). Some of their regular packing is extremely open, and it seems that the comparatively low C values of natural sands can be explained by the presence of only a small proportion of the grains in an open-packing pattern. When spheres are used, one can take into account only sorting or size variation and not orientation together with size, as spheres have no dimensional orientation. The closer regular packing of equal ellipsoids is highly anisotropic, whereas the open packing is perfectly isotropic. Inasmuch as natural sands possess a fabric which is neither perfectly anisotropic nor perfectly isotropic, a combination of both close- and open-grain packings may be expected to be present on a microscopic or a local scale. For these reasons, Allen (1969) explored the different regular packings of ellipsoids using not the triaxial but the prolate spheroids (= ellipsoid of revolution about the long axis). Figures 3-59 and 3-60 illustrate the nine different packing systems. Figure

@@Do3 Plan

PACKING

Elevat ion

PACKING YIII

Plan

Elevation

Elevation

Q9 Pion

Plan

Fig. 3-60. The five open paikings of equal prolate spheroids. (After Allen, 1969, fig. 2, p. 313; courtesy Geol. Mng.)

DIAGENESIS OF SANDSTONES AND COMPACTION 0.8

I

0

2-

0

C 0.6 L

,

Pocking

'm ' m

'

'

1

0

1

b

1

151

C OM1 0.698 0.605 0.555 0.524

LL I-

z

0.454 0.417

W 0 0.4

z

0

u W

I! 0.2

ka L

0

0

0.2

0.4

0.6

AXIAL RATIO, b/o

0.8

1.0

Fig. 3-61. Concentration as a function of axial ratio for the spheroid packings of Figs. 3-59 and 3-60. (After Allen, 1969, fig. 3, p . 314; courtesy Geol. Mag.)

3-61 shows that: (1)C, for example, decreases increasingly rapidly as the axial ratio falls below unity (case V); (2) C declines increasingly rapidly (more so than in case V) as the axial ratio decreases below unity (case VI); (3) C decreases rapidly with the reduction of the axial ratio below unity (case VII); (4) C at first increases with falling axial ratio, but thereafter decreases rapidly (case VIII); and ( 5 ) C at first increases slightly before starting to decrease rapidly as the axial ratio is reduced below unity (case IX). Aside from the regular, well-defined packings presented here, it is not clear how they are to be used to form a model of the concentration of real sedimentary particles arranged in a partially disordered manner. On the basis of theoretical considerations provided by Allen (1969), it seems that on combining the anisotropic rhombohedra1 arrangement of packing IV with the two-dimensional isotropic packing VI, a field that partly represents natural sands is obtained, as shown in Fig. 3-62. As pointed out by Allen (1965, pp. 317-318), only a comparatively small proportion of spheroids in an open, isotropic packing is necessary to give concentrations in the observed range of natural sands. Theoretically, the particles of a natural sand should show a substantial degree of dimensional ordering, though not a perfect one. The degree of ordering should increase with increasing grain concentration, C, for a constant axial ratio. After the theoretical studies, an attempt was made t o compare the results with natural sands, e.g., the grain relationships were investigated. In order to exclude the possibility of secondary changes of the original depositional

K.H. WOLF AND G.V. CHILINGARIAN

152

z-

s I- 0.6 U

a I2 W

0.4

0

u w

-J

I-

0 0.2

a

a!

0

0.4 M 0.8 I.o AXIAL RATIO, b/a Fig. 3-62. Concentration of a spheroid assemblage of mixed packings (IV and VI) as a function of axial ratio and proportion of packing VI given as (13). (After Allen, 1969, fig. 4, p. 317; courtesy Geol. Mag.)

0

Q2

fabric by compaction, for example, the samples collected had relatively high contents of secondarily introduced cement and lacked: (1) empty pore space; (2) evidence of recementation and corrosion of the detrital grains; and (3)evidence of physical distortion or breakage of particles. It was found that the natural sandstones had local patches of grains packed according to the V through IX patterns (Fig. 3-62),which although uncommon were not rare. In the future, maybe statistical analyses can be performed on the occurrence and abundance of these various packing patterns. The results shown in Fig. 3-63are as follows: a,b = long axes virtually at right angles, packing V; c-e = three elongated grains, in a manner as in packings VI and VII; f-i = trios of grains, similar to packings VIII and IX; j , h = two sets of four grains each arranged roughly along the edges of a square or rectangle, similar to packing VI; and 1-p = clusters of a large number of grains arranged in a manner similar to packing VIII. From these results, Allen (1969,p. 320) concluded that the “concentration of a partially disordered assemblage of spheroids can be represented by a model in which one open and one close regular packing are combined in a proportion to yield observed concentration in natural sands and sandstones. In terms of such a model, the concentration of a natural sand could be achieved if a comparatively small percentage of an open regular packing were combined with a comparatively large percentage of a close regular packing. Moreover, the percentage would be such that the sand would stiH display a substantial degree of dimensional ordering of the particles.”

DIAGENESIS OF SANDSTONES AND COMPACTION

153

Fig. 3-63. Cluster of real grains simulating the packings of Fig. 3-60, as observed in sandstones of Old Red Sandstone age. (After Allen, 1969, fig. 5, p. 319; courtesy Geol. Mag. 1

The concept of Morrow Morrow (1971) used a concept that is not always given sufficient consideration, possibly as a result of difficulties in measuring the parameter which he termed *‘packingheterogeneity”. He defined it as ‘‘local variation in sorting or, more strictly, local variation in pore-size distribution”. By employing Figs. 3-64to 3-67,he gave examples of homogeneity-heterogeneity determined by grain-size and packing variations. In Fig. 3-64there is no variation in grain-size distribution from region to region within the illustration, whereas there is such a variation in Fig. 3-65,which makes the texture homogeneous and heterogeneous, respectively. In Fig. 3-66the style of packing is random, whereas in Fig. 3-67part of the particles show cubic and other hexagonal cells, so that again the texture is homogeneous and heterogeneous, respectively. Suction (capillary pressure) required to drain the samples is inversely proportional to the pore size. The slope of the capillary pressure curve obtained in the laboratory experiment (e.g., see Morrow, 1971) re-

154

K.H. WOLF AND G.V. CHILINGARIAN

Fig. 3-64. Homogeneous distribution of rock grains illustrated as monolayer. (After Morrow, 1971,fig. 1, p. 515;courtesy Am. Assoc. Pet. Geologists.) Fig. 3-65.Heterogeneous distribution of rock grains illustrated as monolayer. (After Morrow, 1971,fig. 2, p . 515;courtesy Am. Assoc. Pet. Geologists.)

flects the pore-size distribution of the sediment, because the larger pores tend to drain before the smaller ones in a definite sequence based on pore size. The flatness of the capillary pressure curve, i.e., when the major portion of the curve tends t o be parallel t o the abscissa, is indicative of uniformity in pore size. The curve will show an “irreducible water saturation” (Fig. 3-68A) independent of further increase in the externally applied pressure, so that this property is a definitive, characteristic property of a rock (see Introduction chapter, Vol. 1).According to Morrow, fluid and rock variables (i.e., interfacial tension, viscosity, fluid density, visco-elasticity, particle shape,

Fig. 3-66.Homogeneous (random) packing of spheres illustrated as monolayer. (After Morrow, 1971, fig. 3, p. 516;courtesy Am. Assoc. Pet. Geologists.) Fig. 3-67.Heterogeneous packing of spheres containing high proportion of cubic and hexagonal cells illustrated as monolayer. (After Morrow, fig. 4, p. 516; courtesy Am. Assoc. Pet. Geologists.)

DIAGENESIS OF SANDSTONES AND COMPACTION

155

.ARREOUCIBLE SATURATION

SATURATION, %

Fig. 3-68A. Form of typical capillary pressure drainage curve. Irreducible saturation is dependent upon the pore geometry but independent of absolute pore size, whereas range of displacement pressures is dependent on absolute pore sizes. (After Morrow, 1971, fig. 5, p. 517; courtesy Am. Assoc. Pet. Geologists.) Fig. 3-68B. Retention of liquid at irreducible saturation; rock grains are shown as a monolayer. (After Morrow, 1971, fig. 6, p. 517; courtesy Am. Assoc. Pet. Geologists.)

size and distribution, and porosity) have little influence in themselves on the magnitude of irreducible saturation. Morrow’s experiments demonstrated that irreducible saturation can be an index of packing heterogeneity. He even suggested packing heterogeneity, which can be used as a supplementary characterization parameter of porous rocks, may be found to correlate with depositional environment and changes due to diagenetic processes, such as cementation and compaction. It has been found that irreducible saturation of sedimentary rocks increases as permeability and/or porosity decreases, for example. As to the relationship between irreducible saturation and packing, one should note that a variety of packings can give the same irreducible saturation, so that correlation is not a simple one. Figure 68B depicts the retention of liquid at irreducible saturation. The concept of “packing heterogeneity” must be given full consideration in the study of textures as a result of compaction, because any primary heterogeneities will control the subsequent style of grain movement during compaction.

156

K.H. WOLF AND G.V. CHILINGARIAN

The measuring technique of Kahn Kahn (1956) offered a method to measure packing and stated that packing is the mutual spatial relationship among the grains. Packing can be measured in thin sections by traversing the rock slice using a mechanical stage; the ocular must have a cross-hair and a micrometer to measure the length of the grains traversed (Fig. 3-69). As explained by Kahn and illustrated in Fig. 3-69, two concepts or measurements are used:

(1)Packing proximity, Pp = q/n X 100

(3-1) where q = number of grain-to-graincontacts, n = total number of contacts, as well as total number of grains, and 0 < q < n.

cgilt) n

(2) Packing density, Pd =(m

i=1

X 100

(3-2)

where m = correction term for various combinations of ocular, objective, and scale, t = total length of traverse, g = intercept values, n = the total number n t . of grains in a given traverse across the thin section, and 0 < i = 1 gi rn

c

Packing proximity, Pp, measures property of packing, expressed as the percentage of grain-to-grain contacts in a traverse of n contacts, whereas packing density, Pd, measures the aggregate property of packing. Kahn’s technique has been used with success, e.g., by Martini (1972).

Depth of burial and diagenesis (Taylor, 1950) Taylor (1950) studied several sequences of sandstones to determine what influence depth of burial has had on diagenesis, e.g., as reflected by graincontact varieties. The sandstones investigated by her cemented very early, and their textures and fabrics were controlled by depositional and accumulational factors. On the other hand, when induration occurred after additional beds of sediments accumulated (i.e., after burial), the role of many chemical and/or physical diagenetic factors may have been very important in determining the origin of secondary textures and fabrics. The controlling factors included: (a) mineralogic composition of the sand; (b) roundness and sphericity of grains; (c) sorting and grain-size distribution; (d) stratigraphic distribution of sand, i.e., whether a regional blanket or localized; (e) depth of burial; (f) tectonic or diastrophic processes; (g) temperature; (h) ground-water circulation condition; and (i) geologic time. These variables determine the type of pore-space reduction, which occurs as a result of: (1)pore filling; (2) elastic and plastic deformation of grains; (3)

DIAGENESIS O F SANDSTONES AND COMPACTION

c = c a r b o n a t e cement s = silica cement m =matrix .=void

157

-

ith grain 9,: i t h intercept value t = length of t r a v e r s e

ti

Fig. 3-69.Schematic representation of a sand-size sedimentary rock. The total number of grains in the traverse is nine. Of these nine grains there are nine contacts, i.e., grain 1 is in contact with matrix, grain 2 with a rock fragment, grain 3 with a void, grain 4 with matrix, grain 5 with carbonate cement, grain 6 with silica cement (an overgrowth), grain 7 with silica cement, grain 8 with matrix, and grain 9 with matrix. This results in, but one grain-to-grain contact. Packing proximity is, therefore, equal to 1/9 X 100 = 11.11%. Packing density is the sum of the intercept-size values for the nine grains, g, + g, + g, + g4 + gs + g6 + g7 + ge + gg = zr==tgi, divided by the length of transverse, t, and corrected for magnification. (After Kahn, 1956, fig. 2, p. 391; by permission of The Univ. of Chicago Press, copyright @ 1956 University ofChicago.)

solution and redeposition of dissolved material; and (4) plastic flow of material under pressure. Thus, pore-space reduction and compaction are of physical and/or chemical origin. Taylor found that in sandstones which have undergone early cementation (i.e., pressure effects are absent or at a minimum) there are three types of contacts: tangential, long, and concavo-convex. The shape of the grains and original packing control the grain-to-grain fabric. Tangential contacts (Table 3-XX) occurring as points in the plane of a thin section, have maximum development in deposits that were cemented early by pore infilling. Long contacts (relatively straight lines in the plane of thin section, see Fig. 3-47 and ,Table 3-XX) are more variable, just as the causes and controlling factors of their occurrence. A small edge of a grain or only part of an edge may be involved and thus form only a very short line. If grains are flat or have straight edges, then the nature of the final packing and orientation of the

K.H. WOLF AND G.V. CHILINGARIAN

158

grain accumulations control the proportion of long contacts. A sand with packing that results in maximum porosity has a minimum number of long contacts. With increasing degree of compaction, the number of long contacts increases. Concavo-convex contacts (curved line in the plane of thin section as shown in Fig. 3-47) are rarely found in sand in which the porosity has been reduced by monomineralic pore filling. In this case, the concavoconvex contacts are the result of the shape of the grains involved. The number of contacts per grain is small when early cementation took place. Random sections through haphazardly-packed spheres of uniform size have 0.63 contact/grain. Taylor has found 1.6 contacts/grain to exist in experimental sand deposits. Variations in size and shape are important in determining the number of contacts. Floating grains (no contact in the plane of thin sections as illustrated in Fig. 3-47) occur in sandstones that underwent early cementation. It should be pointed out, however, that in three dimensions these grains are usually supported. Chance packing of spheres gives rise to 47% floating grains, whereas the experimental sand prepared by Taylor had about 17% floating grains. In some cases, the power of crystallization of cement between the grains may have pushed them apart to form the “floating-grain” fabric. This may be indicative of either very early diagenetic precipitation of the cement and/or suggest a lack of thick overburden. This

EXPERIMENT&

SAND

FLOATING -TANGENTIAL=

LONG

CONCAVO-CONVEX- NONE

j

Fig. 3-70. Pressure phenomena, types. of contacts, and nature and percentage of secondary pore filling in Wyoming sands as related to depth. (After Taylor, 1950, fig. 1, p. 7 0 8 ; courtesy Am. Assoc. Pet. Geologists.)

DIAGENESIS OF SANDSTONES AND COMPACTION

159

interpretation applies to the volcanic arenites investigated by Wolf and Ellison (1971), which have a disrupted framework resulting from the calcite cement precipitation. Figure 3-70 indicates the general distribution of cement (<7%) in the sandstones studied by Taylor. In this figure, the dashed line represents less than 1%of the particular grain-contact type under consideration, whereas the full line depicts less than 10%. The various grain-to-grain contacts are also shown in Figs. 3-71 and 3-72. Pore-space reduction by plastic flow of material and by solution-recrystallization (or precipitation) of material occurs under pressure as a response t o burial and, perhaps, is controlled by the increase in geothermal gradient. Evidence of pressure is demonstrated by sandstones in which the volume of pore space plus the volume of chemical cement is less than the original pore volume of the sand. If a well-sorted sand has three or more contacts per grain, pressure probably has reduced the porosity. Pressure is also indicated where (1)a relatively large percentage of the contacts are long, concavoconvex, and sutured, and (2) grains are fractured, crushed, or bent. Although the above three contact types are chiefly due to pressure, concavo-convex and long contacts may also be due t o the nature of original packing and grain shape, as discussed above. Concavo-convex contacts are due to (a) plastic flow of the yielding grain, and (b) solution at points of contacts between grains. In the latter case, the dissolved material may be redeposited in nearby pores or removed in solution. DEPTH AND FORMATION 2885

4535'

' X P ~ ~ ~ ~ESAVERDE T A L SHANNON NUMBER OF CONTACTS PER GRAIN

"FLOAT IN G' GRAINS

CONCAVOCONVEX

SUTURED

4.4

I

4.9

1

5.2

' -

-

.3%

16 64h

I

9.6

~

191

.9

-

-

51.6

51.5

45 0

28 5

28 I

23 I

18 5

I9 7

31.8

Fig. 3-71. Number'and types of contacts in experimental and Wyoming sands. (After Taylor, 1950, fig. 2, p. 7 1 4 ; courtesy Am. Assoc. Pet. Geologists.)

K.H. WOLF AND G.V. CHILINGARIAN

160

+

-

3 60 v -

a

w40n

-

20

0

4 *.......... I

0

I

2000

....... .............. ” , 4.*:.. O-..-o...

3

-

1

......

I

I

4000

6000

8000

DEPTH Fig. 3-72. Graph showing number and type of contacts in Wyoming sands. 1 = tangential contacts; 2 = long contacts; 3 = number of contacts; 4 = concavo-convex contacts; 5 = sutured contacts. (After Taylor, 1950, fig. 3, p. 175; courtesy Am. Assoc. Pet. Geologists.)

Plastic flow has been observed in quartz, feldspar, and chert as well as in rock fragments, especially in shales, mudstones, phyllites, and schists. The numerous factors which determine the tendency of grains to yield include crystallographic orientation, chemical structure and composition, physical properties, original shape, amount and direction of pressure, amount of heat, amounts and types of matrix and interstitial fluids, and relative solubilities of the minerals. Taylor (1950, p. 710) stated that sutured contacts are due to local solution which produces wavy or zig-zag contact lines in the plane of thin sections. Impurities, such as clay, iron oxides, and bituminous material originally present as dust rings on the grains, are left behind as “insolubles” along the contacts. Extensive pressure solution may result in microstylolite contacts. Long contacts are due to (a) original packing of grains that have straight sides or edges, (b) precipitation of secondary cement, and (c) pressure. Consequently, great care must be taken in making an interpretation. In case a, some long contacts may result very early during compaction when grains rotate and adjust themselves. When plastic flow occurs in cases where grains have a straight edge, long contacts form. On the other hand, where adjacent grains are round, concavo-convex contacts will result. Long contacts between grains may represent a transitional stage between the original tangential and final sutured contacts. Figure 3-70shows the number &d types of grain-to-graincontacts, whereas in Fig. 3-72this data is presented graphically. There is a definite increase

DIAGENESIS OF SANDSTONES AND COMPACTION

161

in the number of contacts per grain with increasing depth, from an average of 2.5 contacts/grain in the Mesaverde Formation at 2885 f t to 5.2 contacts/ grain in the Morrison Formation at a depth of 8343 ft, with three intermediate values. The type of contacts shows a progressive change with depth, from those controlled by depositional packing to those resulting from pressure. Each curve in Fig. 3-72 has its own specific trend with depth. No relationship to depth of burial has been established for crushing, flowage, development of twin lamellae, cracking, and bending, which have also been recorded. As Taylor (1950,p. 716) has pointed out, the above-mentioned type of investigation is useful in studies of: (a) porosity and permeability reduction in vertical and horizontal sequences; (b) types and relative importance of processes that cause pore-space changes in sandstones that differ in composition, environment of deposition, and postdepositional history (i.e., diagenesis-burial epigenesis-metamorphism); (c) predicting porosity and permeability at different depths; and (d) determining the movements of hydrocarbons and ore fluids (see approach by Fuchtbauer, 1961, 1967a; and separate section on p. 311). Numerous investigators have employed the technique described by Taylor. Gaither (1953)presented some interesting results of investigations of packing in laboratory-sedimented, medium-grained St. Peter sand. The well-sorted sand with a porosity of 37% showed the following fabric: (a) 0.85 contacts per grain in the thin sections prepared; (b) 46% of the grains were in a floating position; (c) 31% of the grains were in contact with only one other grain; (d) 16% of the grains were in contact with two other grains; (e) 6%of the grains were in contact with three other grains; and (f) 1% of the grains were in contact with four other grains. More than three-quarters of the grain-to-grain contacts were of the tangential type, and most of the rest were long contacts. Concavo-convex contacts were uncommon, whereas sutured contacts were absent. Wright (1964)used Taylor’s (1950)approach in his investigation of clastic rocks. Figure 3-73 shows the results of a series of quantitative determinations of the number of grain contacts. Wright (p. 767) stated: “for most sediments the histogram roughly follows the shape of a probability curve, but in very poorly sorted sediments a tail occurs at one side of the histogram” (No. 15, Fig. 3-73).This tail represents the large grains, which are in contact with a large number of the surrounding smaller grains. All those histograms that have maxima around one grain contact (Nos. 6 and 8, Fig. 3-73),represent rocks that underwent early cementation. Many of the sediments studied contained clay minerals as a matrix that was not solid or resistant enough *toprevent compaction, but prevented compaction to proceed as far as in the clay-free deposits. Wright (1964,p. 758), therefore,

K.H. WOLF AND G.V. CHILINGARIAN

162

0

m ' 0 2 4 6 1

0 2 4

NUMBER OF GRAIN-CONTACTS PER GRAIN Fig. 3-73.Histograms prepared from determinations of the number of grain contacts per grain in fifteen thin sections of sandstones and siltstones. Sections 1-1 0 and 15 = Middle Grit Group of the Black Hill area, Holmfirth; 6 = siltstone below the Readycon Dean Series; 2 and 15 = Readycon Dean Series; 4 = siltstone below the Heyden Rock; 1, 3, 5 and 7 = from the Heyden Rock; 8, 10 and 11 = silts and sandstones above the Heyden Rock; 9 = Huddersfield White Rock; 12 = Rough Rock of Staffordshire; 13 = the Morridge Grit, Staffordshire; 14 = Elland Flags, Fagley Lane Quarry, Bradford. (After Wright, 1964,fig. 1, p. 757;courtesy ?J. Sed. Petrol.)

concluded that some account must be taken of the clay content before the significance of the average number of grain contacts can be assessed (Fig. 3-74),There are two factors which control the average number of grain contacts: depths of burial and proportion of non-quartzitic constituents in the quartz-rich deposits. Whether the number of grain contacts increases with depth in a proportional fashion or not has not been determined, neither by Taylor (1950)nor by Wright (1964).Figure 3-74,however, shows that there is a relationship between average number of grain contacts and percentage of quartz (or the percentage of non-quartz minerals). The small scatter may be due to variation in sorting. The straight line in Fig. 3-74cuts the grain-contact axis at a value of just over six, meaning that if the rocks were purely quartzitic (i.e., no day matrix), an average number of contacts per grain of about 6.2could be expected. Based on Taylor's observation that the

163

DIAGENESIS OF SANDSTONES AND COMPACTION

1

Lu

0

0

I

1

t

CEMENT

1

GRAINS

PERCENTAGE OF NON-QUARTZ CONSTITUENTS

I:) GroinlMatrix Ratio

MATRIX

Fig. 3-74.Modes of grain-contact histograms (Fig. 3-73)plotted against the percentage of non-quartzitic constituents. (After Wright, 1964,fig. 2,p. 758;courtesy J. Sed. Petrol.) Fig. 3-75. Ternary classification of sedimentary rocks using fundamental end members and showing range of samples studied. (After Smith, 1969, fig. 1, p. 261;courtesy Am. Assoc. Pet. Geologists.)

sandstone at a depth of 8300 f t in subsurface had 5.2 grain contacts per grain, an extrapolation was made by Wright for his own rocks, which suggested a burial depth of about 10,000 ft. A crucial question arises, of course, as to what extent extrapolations of this kind are reliable in other instances and an answer remains to be found by future investigators. Overburden pressure alone may not be the only parameter that controls the number of grain contacts and the type of contacts, because possibly the geothermal gradient and pore-solution chemistry, to name only two variables, may be effective also.

Puartz

I

95

75 25

F-Well Cores I

50 50

I

!

25 75

- I _Tyrone I

5 1 Dolomite 95 I ( O h )

Section

I

Fig. 3-76.End member dassification of dolomite-quartz system showing range of samples studied. (After Smith, 1969,fig. 2,p. 263;courtesy Am. Assoc. Pet. Geologists.)

K.H. WOLF AND G.V. CHILINGARIAN

164

Study o n a carbonate-quartz system (Smith, 1969) Smith (1969)investigated the relationships among petrography, porosity, and permeability in a carbonate-quartz system, which is particularly interesting as most investigations by other researchers were performed on monomineralic rock suites. Smith collected his specimens from two sedimentary rock suites of which one set was composed of nearly pure (quartz-poor) dolomite and the other of nearly pure (carbonate-poor) quartzite, as diagrammatically presented in Figs. 3-75 and 3-76.They are referred to as “F-well cores” and “Tyrone section specimens,” respectively. As illustrated in Fig. 3-76,the samples of the F-well cores range in lithology from 40% quartz (= sandy dolomite) to 100% quartz (= orthoquartzite). Figure 3-77shows a plot of porosity versus insoluble residue (I.R.) with a definite positive relationship. A few scattered points represent relatively higher porosity values than would be expected from their corresponding I.R. values. In general, porosity increases with increasing I.R. Smith mentioned several possible factors that may have resulted in this deviation, including compaction and cementation. There are two groups noticeable in Fig. 3-77, as shown, for example, by the relative flatness of the curve below 30--40% I.R. Fig. 3-78 shows the relationship between packing and porosity, which appears to be statistically significant and positively correlative. From the correlation in Fig. 3-77,one might expect a statistically significant positive relation between packing and I.R., which is indeed the case as shown in Fig. 3-79.One should note that between 80 and 100% I.R.,the degree of packing varies over a range of nearly 50%.

3

>:

“1 IS

::

INSOLUBLE RESIDUE, X

Fig. 3-77. Scatter diagram showing relation between percent insoluble residue and percent porosity for F-well ( 1 ) hnd Tyrone (2) samples. (Modified after Smith, 1969, fig. 6, p. 268; courtesy Am. Assoc. Pet. Geologists.)

DIAGENESIS OF SANDSTONES AND COMPACTION LO

1

1

I

I

I

I

1

I

165

I

0 0

0

I6

-

0

0

O

O 0

0

0

0

0

aQ

s 12-

0

5 P

o

2

-

0

0

~

0 0

0 8

0

o

o

! ? -

0

o

0

0

o

0

o

1 0

0

0

0

D

0

0

0 0

4 r

l o

0

0 - O '

0

'

20

I

I

40

60

I

80

I

~

K)o

0

40

80

PACKING, %

Fig. 3-79.Scatter diagram showing relation between percent packing and percent insoluble residue in F-well samples. n = 56; r = 0.489. (After Smith, 1969, fig. 8, p. 269; courtesy Am. Assoc. Pet. Geologista.)

166

K.H. WOLF AND G.V. CHILINGARIAN

Smith concluded that for this carbonate-quartz system an increase in porosity accompanies an increase in I.R. (quartz) between about 50 and 100% I.R. Part of the increase may be the result of changes in the degree of packing of the quartz grains. Intergranular porosity, which certainly is related to the packing (Fig. 3-78),was observed under the microscope in most of the samples containing 75-100% I.R. Based on statistical component analysis, Smith (p. 272) concluded that “the dependent property, porosity, is associated closely with the independent properties, I. R. and packing. The other independent properties, quartz grain size, sorting, shape, and standard deviation of shape, are unrelated to each other, and also are unrelated t o porosity, packing and I.R. The factors controlling the variability in porosity, packing, and I.R. therefore are unlike those controlling size, shape, and sorting, or the latter properties respond differently t o the same sedimentary processes.” “. . . porosity varies primarily as a function of the amount of I.R. and the degree of packing in the samples. There is no evidence to suggest that porosity is influenced appreciably by grain size in these sedimentary rocks, although grain size commonly has been shown to be one of the most important variables influencing porosity in other studies of this nature. Shape and sorting of the quartz grains in these samples also are unrelated to porosity.” In the section on the Framework Concept, Smith (p. 274) stated that the original porosity of the carbonate-quartz system was reduced as compaction continued until the framework of the grains began to support the overburden and thus preserved the remaining pore spaces. In the case of a pure limestone or dolomite composed of non-granular components (i.e., made up of lime mud or dolomite mud, which is more compactible than the deposits of oolites, skeletal fragments and pellets, for example), compaction reduces the pore space considerably more than it does in sediments with granular limy and quartz constituents. Continued addition of clastic or limy granular components to the non-clastic carbonate phase will eventually result in a sediment composed of enough grains in contact with each other to form a supporting framework. According to Smith, this type of granular framework will reduce or prevent compaction and generally preserve the larger pores in the sediment. As to a deposit composed of quartz grains that make up the framework, with a carbonate cement filling some or all of the intergranular pores, Smith (p. 275) stated that the final porosity in this type of carbonatequartz system depends on: (1)the original intergranular porosity which is a function of packing, for example; (2) the amount of secondary cement; and (3) the amount of cement removed by subsequent decementation. Smith compared the results of his work with those of Lucia (1962)and concluded that dolomitization of th*e rocks investigated by him occurred before compaction.

DIAGENESIS OF SANDSTONES AND COMPACTION

167

Complex diagenetic alterations (Aalto, 1972)

Aalto (1972)presented a fine example illustrating the complex interrelationships between textural, compositional and deformational factors for an orthoquartzitic or orthoquartzitic-sub-graywacke suite, with commonly uniform grain composition and complex diagenetic alterations among grains, cements, and clay matrices. In order to understand the terminology given in Figs. 3-80 and 3-81,the following information on the six grain types is necessary: (a) Undulose quartz (0-25" rotation) often has quartz overgrowths. Many grains contain bubble-like gas or liquid inclusions arranged in linear trains. In very deformed grains, these inclusions form parallel, closely-spaced trains

$ N ka 3 0 W

c

I

a

PHI MEAN SIZE

C

80 60

40

20

b

d

DEPOSITIONAL MATRIX (YO)

Fig. 3-80. Interrelationships of textural and deformational parameters. (a) Percentage of polycrystalline quartz present versus phi mean size; (b) percent of grains with overgrowths versus estimated percent of depositional matrix; (c) percent of grains with linear trains of inclusions and comb-like groups (clg) of trains versus phi mean size; and (d) percent of grains wit% linear trains of inclusions and comb-like groups of trains versus estimated percent of depositional matrix. (After Aalto, 1972, fig. 9; courtesy J. Sed. Pe tro 1. )

168

K.H. WOLF AND G.V. CHILINGARIAN CONCAVO-CONVEXlccl and SIMPLE LlNEIlbl

Fig. 3-81.Types of grain contacts. Letters presented here represent the samples studied and correspond to particular formations as follows: Q = Gutierrez-Quetame Sandstone; 2 = Cdqueza Sandstone; U = Une Sandstone; P = fine sandstones and siltstones of the lower part of the Guadalupe Group; D = Dura Sandstone; L = Labor Sandstone; T = Tierna Sandstone; G = La Guia Sandstone; C = El Cacho Sandstone; R = La Regadera Sandstone. The summary diagram (small triangle on left) portrays general relationships among textural, compositional and deformational parameters. (After Aalto, 1972, fig. 11, p. 338; courtesy J. Sed. Petrol.)

and resemble the teeth of a comb. Inter- and intragranular fractures are present in all sandstones studied by Aalto, but are most common in the oldest and most deformed rocks. Healing of these fractures by quartz gave rise to the linear trains of “comb-like” groups of bubbles. Microlites of micas, rutile, sphene, tourmaline and zircon are present in the quartz grains also.

DIAGENESIS OF SANDSTONES AND COMPACTION

169

(b) Semi-composite quartz grains have undulose extinction and overgrowths. The overgrowths and the deformational features are similar to those found in monocrystalline quartz, but are less common. (c) Composite quartz grains have undulose extinction, commonly sweeping over the entire grain. Overgrowths are uncommon, but where present they develop in optical continuity with different grains within the composite fragment. The inclusion trains and comb-like groups of trains are confined to individual grains. (d) Stretched quartz has many subparallel, elongate quartz members, and the boundaries are intensely crenulated, sutured or granulated. Strong undulose extinction commonly sweeps the entire grain, whereas the overgrowths, inclusion trains, and comb-like groups of trains are uncommon and only present in individual grains. (e) Polycrystalline quartz grains are present also. (f) Chert grains may be recrystallized, They may contain microfossils, glauconite, and silt and mud grains, all of which may be deformed and altered. Heavy minerals present as traces may show cleavage. (g) The matrix and cement is largely of diagenetic origin with kaolinite, illite, and sericite being common in older sandstones. Details on the diagenesis of “matrix-cement” and alterations of “grain-matrix-cement” and “grain-grain” are presented by Aalto. The interrelationships established by him among textural, compositional and deformational factors are shown in Figs. 3-80and 3-81.As pointed out by Aalto (p. 338), generally poor correlations exhibited in Figs. 3-79b,c, and d , and Fig. 3-81 are due to the presence of deformational features in grains formed prior to deposition by complex postdepositional deformation and by obliteration of grains by iron oxides during epidiagenesis. The data distribution in Fig. 3-81 suggest that sandstones characterized by a high percentage of simple irregular penetration contacts, microstylolite contacts, and pseudograins are commonly coarse-grained. These sandstones also contain more polycrystalline quartz, grains with inclusion trains, and grains with comb-like groups of trains. Aalto (p. 338) also stated: “Moreover, these sandstones have fewer floating grains and a clay or sericitic matrix. In contrast, sandstones with an abundance of simple line and concavo-convex contacts are more likely to be fine-grained and to have less polycrystalline quartz.” They also contain fewer grains with inclusion trains and comb-like groups of trains, more floating grains, and a greater variety of matrices and cements. Inasmuch as crystal face contacts are a product of overgrowth formation, sandstones with many such contacts commonly have more overgrowths. As shown in Fig. 3-80,with an increase in the amount of depositional matrix, sandstones commonly show a decrease in abundance of overgrowths on grains. With increasing amounts of depositional matrix, fine-

K.H. WOLF AND G.V. CHILINGARIAN

170

grained sandstones correspondingly have fewer grains with inclusion trains and comb-like groups of trains (Figs. 3-80 c , d ) . Aalto (p. 339) pointed out that exceptions have been observed for each one of the above generalizations, and that the observations described above reflect a complex interaction of several variables. Irregular and microstylolitic contacts, pseudograins, inclusion trains, and comb-like groups of trains (see Whisonant, 1970), associated with strong undulosity, all reflect increases in deformational stresses and presence of special chemical environment. Porosity and packing index Based on data presented by Masson (1951), Griffiths (1967b) related porosity to packing index (see definition on p. 135), as shown in Fig. 3-82. The artificially-packed sand had an index of 10, whereas the packing index of 32 sandstones from the Gulf Coast and U.S.A. Midwest oil fields varied from 18 to 62. Rosenfeld (1953, in Griffiths, 196713) using Masson’s technique, found that the average packing index of 6 1 orthoquartzite specimens (Fig. 3-83) was 41.4%, with a standard deviation of 17.7% and a range of 3 to 75%. Effects of textures and composition (Renton et al., 1969) Renton et al. (1969) studied the effects of textures and composition on compaction for various types of sands. One set of experiments was performed on quartz crystals surrounded by small grains but with no liquid, in order to simulate purely mechanical effects. After compaction, the quartz crystals showed visible fractures at points of contact and both radial and

0

PACKING INDEX

Fig. 3-82. Relation of porosity to packing index (After P.H. Masson, 1951, in Griffiths, 1967b, fig. 8-2, p. 1 6 9 ; copyright @ 1967, McGraw-Hill, New York.)

DIAGENESIS OF SANDSTONES AND COMPACTION

171

ui 12 w

J LL

5r n 8 LL

0

K :4

5z

0

0

20

40

60

PACK I NG I N D E X,

80

O '/

Fig. 3-83.Frequency histogram of grain-to-grain contacts in 61 specimens of Oriskany quartzite, f (mm arithmetic mean) = 41.4%; 8 (standard deviation) = 17.7%. (After Rosenfeld, 1953, in Griffiths, 1967b, fig. 8-3,p. 169, copyright 0 1967 McGraw-Hill, New York.)

TABLE 3-XXII Compaction of quartz and chert sands (after Renton et al., 1969,table 3, p. 1113) Expt.

Maximum Duraload (psi) tion (days)

Experimental material Initial Final Corn% void poros- poros- paction decrease ity(%) ity(%) (%)

7

6000

48

16-21 mesh round

6

6000

48

8

6000

47

14

6000

39

10

4000

23

11

4000

23

No.

quartz grains 60-80 mesh round quartz grains 60-80 mesh angular quartz grains angular quartzsilt (0.04mm) 60-80 mesh angular chert grains mixture 60-80 mesh angular quartz (50%) and angular chert (50%)

36.5

24.5

16.2

44.0

39.0

16.5

28.0

69.5

45.8

13.0

37.8

81.5

53.5

17.0

43.5

80.3

49.3

15.9

39.3

80.5

51.0

27.0

32.7

64.5

All experiments were conducted at 400"C with 0.5M NaZC03 solutions at a hydrostatic pressure of approximately 6000 psi. In experiments on quartz, load was padually increased to the maximum value over a period of 4 weeks and held at this value for the remainder of the experiment. Load gradually increased during the entire experiments on chert.

K.H. WOLF AND G.Y. CHILINGARIAN

172

annular fractures were present. The radial fractures were perpendicular to the surface of the crystals, whereas the annular ones formed nested cones. In the absence of intergranular liquid, which would have caused dissolution of the quartz at the points of stress, the quartz lattice failed as a brittle substance. Distinct pressure-solution pits were formed in distilled water, NaOH, Na2C03, and NaCl solutions, and natural brines. In a series of experiments, the relative rates of pressure solution and cementation in sands of different types were determined. The porosity, degree of compaction, and percent void reduction were measured and are given in Figs. 3-84 to 3-87. Porosity values before and after the compaction are presented in Table 3-XXII. According to Renton et al., compaction was predominantly the result of pressure solution at grain contacts, because no appreciable fracturing occurred and many concavo-convex contacts formed between grains. SiOz was precipitated as euhedral quartz overgrowths on the grains.

DURATION, days Fig. 3-84. Porosity reduction in various angular quartz and chert sands under t h e same conditions of compaction for each sample. On Figs. 3-84 to 3-87, solid symbols represent values based o n actual volume measurements, whereas open symbols represent values calculated from t h e amount of fluid removed during the experiments. As in Fig. 3-85, 5 = chert-quartz mixture; 4 = angular quartz silt (0.04 mm); 3 = 60-80 mesh angular quartz; 6 = 6--80 mesh chert. ( A f t e r Renton et al., 1969. Fig. 10,p. 1112; courtesy J. Sed. Petrol.)

DIAGENESIS OF SANDSTONES AND COMPACTION

173

Figure 3-86demonstrates the following facts: (a) The compaction rates decreased with time as the areas of contact surfaces of the grains increased and the load was distributed over larger areas. (b) Finer sands underwent greater compaction with resulting greater reduction in porosity, probably due to larger numbers of contact points. (c) Compaction of chert grains was more rapid than that of monocrystalline quartz. This may be attributed to the small crystal size of the chert or the presence of poorly-crystallized material very closely resembling a homogeneous chert, which facilitated more rapid pressure solution. (d) In mixtures of 50%chert and 50% quartz fragments, the initial com-

W-

v)

Q

W

a

50

0

I0

20

30

DURATION, doyr

40

50

0

10

20

30

40

50

DURATION, day8

Fig. 3-85.Variation in compaction rates of different types of sands under the same loading conditions that remained unchanged from sample to sample. 1 = 16-21 mesh round quartz; 2 = 60-80 mesh round quartz; 3 = 60-80 mesh angular quartz; 4 = angular quartz silt (0.04 mm); 5 = chert-quartz mixture; 6 = 60-80 mesh chert. (After Renton et al., 1969,fig. 12,p. 1113;courtesyJ. Sed. Petrol.) Fig. 3-86.Void reduction in sands of different types under the same experimental conditions that did not change from sample to sample. Legend as in Fig. 3-85.(After Renton et al., 1969,fig. 13,p. 1114;courtesy J. Sed. Petrol.)

K.H. WOLF AND G.V. CHILINGARIAN

174

paction occurred about as fast as in pure chert samples. The rate was similar to that of quartz, however, during the late stages of compaction. As shown in Figs. 3-84and 3-87, in the early stages of compaction, the pressure solution affected chert first. When the quartz grains were largely in contact, the rate of pressure solution was considerably less. Only 7% of the sample was composed of chert after 33%compaction. Another conclusion reached by Renton et al. (p. 1116) was that the rate of compaction and porosity reduction from pressure solution varies considerably depending on size, shape and type of sand grains (see Figs. 3-84 to 3-86). Pressure-solution compaction in samples having angular grains was considerably faster than in samples having round grains of the same size. Renton et al. (p. 1116) observed that the appearance of the angular grains after pressure solution and secondary quartz growth was not greatly different from that of round grains after comparable pressure solution. Thus, very little can be deduced in natural sands about the original grain shape after moderate pressure solution, unless remnants of clearly-defined dust rings are preserved.

DURATION, days Fig. 3-87. Porosity reduction in coarse-grained and fine-grained sand under the same conditions of compaction from sample to sample. Legend as in Fig. 3-86. (After Renton et al., 1969, fig. 11, p. 1113; courtesy J. Sed. Petrol.)

DIAGENESIS OF SANDSTONES AND COMPACTION

175

As shown in Fig. 3-86, although the fine angular sand grains had considerably greater initial porosity, their porosity became less than that of the fine round sands after about 25% volume compaction. The contrast between the compactability of angular silt and coarse sand (16-21 mesh) is striking. As pointed out by Renton et al. (p. 1116), the compaction of the silt was over 3 times greater than that of the coarse sand after the same length of time, i.e., 80%and 38%loss (by volume) of voids for silt and coarser sand, respectively. Porosity loss of samples of chert grains was considerably more rapid than that for monocrystalline quartz (Fig. 3-84). This diagram also shows (cf. Fig. 3-86) that there was a marked reduction in the proportion of chert, because in mixtures of chert and quartz grains much of the presolved chert was precipitated as overgrowths on the quartz particles. Hence, original chert contents in natural, compacted sands might be impossible to determine. From the above data, Renton et al. suggested that it is possible to “estimate porosity variation in certain natural sands where the degree of pressure solution was reasonably uniform and the dissolved silica was deposited locally. If the porosity reduction is known for a sand similar to one of the types shown in Figs. 3-84 and 3-85, the reduction can be estimated for the other types given. For instance, if a porosity reduction of 23% has occurred in 60-80-mesh round sand, only a 12% reduction would be expected in 16-21mesh sand. Exact porosity predictions are not possible in natural sands because of the large number of variables involved; however, these data should be of aid in estimating the relative magnitude of the porosity reduction’ from pressure solution.”

Observations on rims and coatings (Pittman and Lumsden, 1968) Pittman and Lumsden (1968) made observations, which were similar to those by Fuchtbauer et al. (see pp. 343-363), indicating that authigenic, fibrous chlorite rims on quartz sand grains inhibited pressure solution. They observed the following types of occurrences: (a) Thick continuous coatings of chlorite which preserve porosity, because pressure solution is retarded and quartz overgrowths are absent. (b) Sparse, thin, and discontinuous chlorite coatings; the sandstones are tight (= non-porous), because pressure solution and quartz overgrowth took place. (c) An intermediate stage between a and b where chlorite coating is thick and complete enough t o have retarded pressure solution and yet is discontinuous in places to allow some local quartz precipitation. Studies such as these have been used in petroleum reservoir investigations and may find application in th’e future in unravelling the origin of certain types of ores in sediments and pyroclastics.

K.H. WOLF AND G.V. CHILINGARIAN

176

Study on the source of silica (Waugh, 1970) Waugh (1970)studied the source of silica, which forms quartz cement in continental red-bed sandstones deposited as barchan dunes in hot and arid desert environments. The cement is present as perfectly-formed bipyramidal quartz crystals. Waugh considered the more commonly proposed sources, including pressure solution, but found that none can adequately explain the origin of the secondary quartz in the sandstones he studied. By establishing that: (a) the number of contacts per grain averages approximately 1.8, (b) over 90% of grains have less than four contacts per grain, and (c) over 90% of the types of grain-to-grain contacts are either floating, tangential or long (Fig. 3-88),he demonstrated that pressure solution did not supply the silica. Waugh also discussed the possibility that silica dust, which is the result of desert abrasion, could have been the source for the quartz cement. Study on grain contacts and packing parameters (Martini, 1972; Gaither, 1953,and others) According to Martini (1972),the effects of compaction are revealed by the number and types of grain contacts and measurements of packing parameters, which describe the threedimensional distribution of the sedimentary particles. This data enables the estimation of the original porosity and permeability of the rock and assists in the interpretation of the secondary changes after deposition. In his investigation, Martini used the packing density, Pd (Kahn, 1956),and packing proximity, P p , as based on grain-to-grain contacts (Taylor, 1950). These packing parameters were measured in the hematitic quartzite to determine possible relationships between them and

A A

A) NUMBER OF CONTACTS PER GRAIN

8) TYPE

FLOATING 8 I

TANGENTIEL

.../

*:\ ,$A/

283

OF GRAIN CONTACTS

4-4’

48’4

LONG

CONCAVO-CONVEX 8 SUTURED

Fig. 3-88. Grain contact analysis illustrating lack of pressure solution effects in the Penrith Sandstone (Lower Permian) of north west England. (After Waugh, 1970, fig. 10,p. 1236; courtesy J. Sed. Petrol.)

DIAGENESIS OF SANDSTONES AND COMPACTION

177

how they relate to other sedimentary properties, e.g., grain size and grain orientation. The results indicated that degree and type of packing in this case correlated strongly with the grain-size variations and that pressure solution (as determined by grain-to-grain contact types) cannot wholly explain the presence of secondary silica cement in the sandstone samples. Inasmuch as the measured distributions of the two packing parameters deviated only slightly from a normal distribution, Martini used a normal statistical testing procedure in establishing relationship between these parameters. As shown in Fig. 3-89,a significant linear correlation is present at the traverse level between packing density and packing proximity and only 10.5% of the variability (scatter) is accounted for by the regression line. Based on the results obtained from the multiple regression analysis, Martini (1972,p. 418) found that better prediction of packing proximity can be made if the intercept size is not considered at the thin-section level. Although a statistical correlation does not imply a causal relationship, the results suggested that as grain size (i.e., intercept size) increases, the amount of space occupied by the particles in a given sample (Pd)tends to increase. Martini (p. 415) stated that “packing proximity was positively correlated with packing density at the traverse level of analysis (five traverses per sample), and that it was best predicted if only the average values of packing density of the samples were considered in a multiple regression analysis”. In relating packing and grain orientation (see Martini, p. 420, for method used), two major relationships were found among packing proximity, packing density, and grain orientation: (1)In sections perpendicular t o the bedding (for grain-imbrication determination) and in direction parallel to the vector mean of grain orientation, a

PACKING DENSITY, f

Fig. 3-89. Average relationships between packing density and packing proximity (circles = average per sample; triangles = average per stratigraphic section; solid line = regression at the traverse level; dotted line = regression at the thin-section level; dashed line = regression at the stratigraflhic section level; square = grand mean). (After Martini, 1972, fig. 2, p. 419; courtesy Int. Geol. Congr., Montreal.)

K.H. WOLF AND G.V. CHILINGARIAN

178

20

' 7 50

60

ro

a

PACKING DENSITY, % ,

PACKING PROXIMITY, %

Fig. 3-90. Relationships among packing density, packing proximity and imbrication of vertical thin sections cut parallel to the vector mean 6 of the grain orientations of the horizontal slides. Significant regression line: Pd = 70.608-1.365 8; r% = 88.67. (After Martini, 1972, fig. 3, p. 420; courtesy Int. Geol. Congr., Montreal.)

significant negative correlation was found to exist between the angle of imbrication and the packing density (Fig. 3-9042). Similar relationship existed between the packing proximity and angle of imbrication (Fig. 3-90,b). The relationships existed in the case of a more open-textured fabric when less space is occupied by the grains. This fabric is due to a higher angle of imbrication, because when the angle of imbrication is higher, relatively coarse grains obtain a better stable, static equilibrium. Also, when a relatively smooth lamina is formed, only the clastic particles with a high imbrication fabric may come t o a stop and be preserved in the microenvironment of sedimentation (Martini, 1972, p. 420). (2) Martini also fuund a significant negative correlation between the percent of vector magnitude of the grain orientation and the packing proximity in cases where P p is less or equal to 18%, whereas for higher values the two variables are independent of each other (Fig. 3-91). Similarly, as also shown in Fig. 3-91, the percentage of vector magnitude of the grain orientation is independent of the packing density. Martini concluded that the correlation between packing proximity and orientation implies that at high Pp values (i.e., when frequent interactions may occur between grains), the degree of preferred grain alignment is independent of the number of contacts between particles. On the other hand, when the Pp values are low (less than 18%),a better degree of grain orientation exists in the presence of fewer grain contacts, i.e., when fewer interactions took place. The grains, which are more elongated and streamlined, have their axes best aligned, and are better oriented in the direction of flow, have a higher probability of stopping when a smooth bed (i.e., with low roughness index) is being formed. Good static equilibrium requires also stronger upflow imbrications of the

DIAGENESIS OF SANDSTONES AND COMPACTION

179

PACKING DENSITY, 7.' 50

40

60

70

w ?

-I 1O 0

20

30

10

15

:

o

.

o

18

20

0 0

25

30

PACKING PROXIMITY, f

Fig. 3-91. Relationships among packing density, packing proximity and vector magnitude per cent ( L )of the grain orientation. Significant regression line: Pp = 15.50-0.279 L ; r2% = 54.36. (After Martini, 1972, fig. 4, p. 4 2 1 ; courtesy Int. Geol. Congr., Montreal.)

particles. In naturally-accumulated sediments, therefore, the sand grains have a preferred orientation and imbrication. The erosion of these grains is more difficult than in instances where the particles are randomly oriented. Martini (1971) did not find a vertical variation in packing parameters in the 40-50 f t thick sandstone sequence, but noticed significant differences on a regional scale. Gaither (1953) presented the results of the investigations of several petrologists who have measured in thin sections the number and types of grain contacts and the porosity of natural sands of varying composition, texture, and postdepositional history. According to him (p. 182), Manry (1949) determined that compaction had reduced the pore space from an assumed original porosity of 38% to an average of 23% in a number of Paleozoic orthoquartzites. Simple pore filling by chemical precipitation of cement resulted in an additional reduction of an average of 7% porosity. In these specimens, the number of contacts per grain ranged from 1.2 to 3.1, which is considerably below the maximum value of 5.2 of Taylor (1950). This may be attributed to: (1)the different lithologies involved; (2) relatively early lithification of the quartzites with a consequent arresting of pore-space reduction; or (3) shallower burial of the quartzites. Beaudry (1950) investigated Pennsylvanian sandstones from depths ranging from 8353 to 13,096 ft, and found that the original porosity of about 37% was reduced to less than 9%. About 60-65% of this reduction in porosity was the result of simple pore filling by quartz and carbonate cement. The remaining 35--40% of pore space reduction was caused by: (1)

180

K.H. WOLF AND G.V. CHILINGAREAN

physical rearrangement of grains; (2) crushing and bending of grains; (3) solution and plastic flow; and (4)replacement and recrystallization. Low average number of contacts per grain (= 2.1-3.6), as compared with those reported by Taylor, and relatively small effects of overburden pressure, in general, were attributed by Beaudry to (1)more resistant composition, i.e., high percentage of quartz grains and low percentage of rock fragments and feldspar, and (2) early introduction of cement. There was no relationship between depth and (a) porosity, (b) number of contacts per grain, or (c) the type of contacts. There was, however, a relationship between the number of contacts per grain and composition, namely, the samples with the most contacts per grain had the largest percentage of calcite grains, whereas those with the fewest grain contacts per grain had the largest percentage of quartz. Hays (1951) studied samples from a depth'of 13,000-21,000 f t and observed that 43-75% of the original porosity was accounted for by simple pore filling, whereas the remainder was eliminated by pressure effects. Average number of contacts per grain ranged from 1.7 to 2.2. Hays concluded that in his samples the number and types of contacts were not related to depth. Samples with the largest number of carbonate grains had abnormally high percentages of sutured contacts. It seems then that, in some cases, composition may have a more important effect on the number and types of contacts than depth of burial. Figure 3-92 shows that when sands prepared in the laboratory are compared with natural, well-indurated sandstones, the percentages of grains with given numbers of contacts, as measured in polished and thin sections, widely

NUMBER OF CONTACTS

Fig. 3-92. Comparison of percentages of grains with given number of contacts as obtained from thin and polished sections of experimental and natural sands. 1 = thin sections; 2 = polished sections; 3 = implied-values for uncompacted sand; 4 = average of Beaudry's (1950) deeply buried sands (8,000-13,000 ft). (After Gaither, 1953, fig. 6, p. 191; courtesy J. sed. Petrol. )

DIAGENESIS OF SANDSTONES AND COMPACTION

181

diverge. In this diagram, the dotted curve 4 represents the distribution of grains with given number of contacts for Beaudry's (1950) deeply buried (8,000-13,000 ft) sandstones. According t o Gaither (1953),p. 193), compaction causes this curve t o differ considerably from the same kind of curve prepared for uncompacted material (heavy broken curve 3). The following four measurable variables can be employed to determine the presence or magnitude of compaction: (1)porosity; (2)average number of contacts per grain; (3) percentage of grains with given numbers of contacts; and (4) percentage of contact types. The first three parameters are closely related and variations in one will be attended by variations in the other two. The most important factors that can cause such variations in these parameters are compaction and sorting. Consequently, interrelationship between compaction and sorting must be thoroughly studied. In the study of undulatory quartz as related to deformation during burial,

q Y z

u

Fig. 3-93. Variation of the undulatory extinction of quartz in relation to the radiogenic age of different granitic rocks. (After De Hills and Corralin, 1964, fig. 1, p. 364; courtesy Geol. SOC.Am. Bull.)

K.H. WOLF AND G.V. CHILINGARIAN

182 100

8 ,-80

-

I

I

I

3 60 0

a

. 92 ..

*. :s

0

>

.

~

H

0

-

t

4 2 20 2

0

3

I

0

0

'

;40 -2, 0

I

,Vo

-

Ia a

I

I

I

-

FAULTED STRATA NON-FAULTED STRATA I

I

I

1

d

the following factors have to be considered: (a) the derivation of undulatory quartz from the source rocks, e.g., granites and metamorphics; (b) differential transportation of undulatory versus non-undulatory quartz (see Conolly, 1965); (c) formation of undulatory quartz during burial; (d) formation of undulatory quartz by tectonism; and (e) conversion of undulatory quartz through recrystallization to less undulatory and/or non-undulatory quartz. By studying granitic rocks of three different ages, De Hills and Corvalh (1964) showed that the intensity of undulatory extinction increases with age as a result of increasing degree of tectonic deformation by successive orogenic movements. Their results are presented in Fig. 3-93. The publications by Blatt (1967) and Blatt and Christie (1963) contain information on properties of quartz from different source rocks, its undulatory extinction, deformation of quartz after deposition, and other related data. Conolly (1965) demonstrated that the percentage of undulatory quartz is greatest near faulted zones, which may also be true in folded rocks (Fig. 3-94). The influence of physical compaction Little information is available on the influence of physical compaction on surface features of grains, and considerable research work is required. It has

DIAGENESIS OF SANDSTONES AND COMPACTION

183

been also noted that certain chemical processes can form “frosting”, for example. These processes may or may not be related to the compaction fluids moving into and out of the sedimentary rock, and it may be impossible to determine the precise origin of the solutions that caused frosting. Surface alterations of grains occurring during calcite cementation or calcite replacement may result in textures that are easily confused with the frosting of eolian sand grains and, consequently, may lead to wrong paleoenvironmental interpretations. Krinsley and Donahue (1968), using an electron microscope, recognized four types of surface textures produced by the microenvironment during diagenesis: crystal surfaces, solution surfaces, pressure-solution striations, and fracture surfaces. One should note that these features are superimposed on surface features which were the result of the original sedimentary and depositional environment. Such studies are necessary t o determine eventually to what extent compaction gives rise t o different surface textures that can be distinguished from each other as to genesis. Margolis (1968) found that (1) sand grains from beaches of low-wave activity were almost all completely covered with crystallographically-oriented etch features; (2) sand grains from littoral zones with moderate wave activity show a combination of mechanical abrasion features and chemical etch features; and (3) quartz grains from beaches with high wave activity predominantly show impact V’s, breakage blocks, and scratches, with very few chemical etch triangles or rhombs. These findings are summarized in Fig. 3-95. Margolis (p. 255) stated that more investigations in other localities and in older sedimentary rocks are required to determine whether the abovedescribed findings d e generally applicable or not. Obviously, they could be applicable only in cases where diagenesis has not obliterated the original features. Margolis (p. 248) remarked: “The degree and the expression of this diagenetic pattern is a function of the length of time the grain has been exposed t o circulating meteoric and ground waters. Other factors that must be taken into consideration are climate, chemical composition of the local ground waters, pH and silica saturation of the waters, and the permeability of the sediment where the sand is located. A high clay content may protect a sand grain from diagenetic change and preserve the original surface textures.” Diagenesis in older sediments may have had a complex history, because the chemical composition and pH of interstitial fluids on the continents can vary greatly with time, from fresh ground water t o connate water during burial and then back to ground water after uplift. Hence, chemical corrosion of sand grains during burial, by compaction or other subsurface fluids, can either obliterate or destroy primary surface textures as mentioned by Margo-

184

K.H. WOLF AND G.V. CHILINGARIAN

v)

W IX

3

2

LOW

Energyl--Moderote I

1

Energy-I-High I

I

Energy I

-

I

50-

-

MEAN WAVE HEIGHT, cm Fig. 3-95. Correlation between mean wave height of the beaches sampled and the occurrence of mechanical features, chemical features, and combinations of both types of features on the surfaces of the sand grains. The ten beach samples were plotted according to the mean wave height (cf. table 1 of Margolis’ paper). The vertical axis indicates the number of grains, out of the total of 50, which exhibited the indicated features. (Energy classification of the beaches from Tanner, 1960). (After Margolis, 1968,fig. 10; courtesy Sed. Geol.)

lis (1968).It may also form new textures that can be easily misinterpreted as being of primary origin. Formation of ‘ ~ l y p t o m o r p h s ” .Rukhin (1958)stated’that the decrease in thickness of sediments during consolidation and compaction is favorable for the formation of “glyptomorphs”, i.e., crystalline aggregates of low solubility that can grow above the sediment’s upper surface. They can be composed of crystals of halite and gypsum, for example, forming within the sediments. Sometimes, these crystalline aggregates include particles of the sedimentary framework. Sand crystals formed by quick crystallization from solution constitute one well-known example. Due to the decrease in thickness of the sediments during compaction, the outline of these crystals may be impressed on the overlying and underlying sediments (Fig. 3-96).

Transitional stages. In the above discussion of sandstone textures and fabrics formed by compaction, no consideration was given t o those related to transitional stages between diagenesis and metamorphism, so that a separate, brief section is devoted to this particular subject. Skolnick (1965), for example, discussed the “quartzite problem” and

DIAGENESIS OF SANDSTONES AND COMPACTION

185

Fig. 3-96. Schematic sketch of the formation of preserved glyptomorphs or casts of halite crystals during compaction o f sediments. A = origin of crystals in newly-formed sediment; B = compaction of the sediment; and C = protruding glyptomorphs above bedding. (After Rukhin, 1958, fig. 56, p. 235.)

mentioned that many sedimentary quartzites are not true “orthoquartzites” (Len, quartz grains cemented by quartz infilling of original intergranular spaces), but have undergone pressure solution during compaction (= “pressolved quartzites”). The latter types are by far the most common quartzites and have certain characteristics that occur also in “metaquartzites”, i.e., formed or modified by metamorphism. Yet, little is known on the effects of time, depth of burial, tectonism, and the properties of the intrastratal fluids on the transition of quartz sand or sandstone to pressolved quartzite and, finally, to metaquartzite. Skolnick’s publication is essential for petrologists, because he described the origin of the ortho- and metaquartzites first and then outlined the terminology and classification of quartzites. In detailed work, it is necessary to distinguish among the three quartz-rich sandstones (i.e., pressolved, orthoand metaquartzites) with numerous possible transitions in between by using the types and frequency of grain contacts. Grain-contact types range from the original depositional varieties to those formed by different stages of compaction and, then, to those that resulted from metamorphism. Com-

186

K.H. WOLF AND G.V. CHILINGARIAN

parative studies of the textures of sandstones changing with depth in sedimentary basins may constitute a promising approach. During the past few years, data on the precise temperature and pressure conditions that cause changes in clays and coaly matter have been collected (for details see pp. 395-422). Thus, in a basin with a variable lithology, the types of quartzites could be compared with the variation in the types and amounts of various clay minerals and, possibly, organic matter with increasing depth. In the Precambrian and Phanerozoic metamorphic belts, the degree and type of metamorphism of clayey sediments and/or volcanics could be compared with the quartzite varieties. Whisonant (1970) reported on slightly metamorphosed sandstones and described postdepositional deformation of quartz-bearing sandstones that can result in the following changes: (a) straining of grains; (b) suturing of grain contacts; (c) solution and redeposition of silica as overgrowths and cement; (d) corrosion and replacement of quartz by sericite; (e) creation of microfaults; (f) formation of fractures either within or transecting the grains; and (g) creation of bubble-trains in healed fractures. The Cambrian rocks studied by Whisonant have undergone early stages of metamorphism as evidenced by secondary sericite, chlorite matrix, straining, suturing, crushing of quartz grains, abundant micro-cataclastic features, microfaults, and fractures. Whisonant (p. 1023) made the following observations on the effects of deformation: (1)As matrix quantity increases, the amount of overgrowth and cementation (by quartz precipitation) decreases; (2) as matrix quantity increases, degree of suturing of grain boundaries decreases; (3) as matrix quantity increases, the amount of straining (as shown by the degree of undulose extinction), which depends on whether the grain is single, semi-composite, or composite, decreases; (4)within the same sample, as grain size decreases, the amount of straining decreases; (5) when matrix is composed predominantly of sericite, as matrix quantity increases, sericitization of detrital quartz and K-feldspar increases; (6) as matrix quantity increases, the number of bubble-trains (i.e., healing microfractures) decreases. In regard to item 4 above, it can be said (see also Conolly, 1965, p. 126) that straining of quartz is a function of grain size and depends on sorting. Where both large and small clasts occur, the finer particles tend to be less strained than the coarser ones, whereas if only fine-grained particles are present, they show no significant difference in the amount of straining. This is due t o larger grains acting as “propsyyin the sandstone during stress application that results in strain shadows in the coarser fragments.

DIAGENESIS OF SANDSTONES AND COMPACTION

187

Whisonant also discussed the derivation of grains, with differing degrees of straining from a source rock area, which were subsequently differentially abraded and dissolved during transportation from the source to the depositional environment. In such cases, the final deposit would be composed of strained grains that differ in both proportion and degree of straining from the source rock material, as a result of the differential processes. Studies like the one mentioned above, will eventually lead to a better understanding of the vague transitional boundaries between diagenesis and catagenesis as well as between catagenesis and metamorphism. Maybe it is possible to make a number of precise subdivisions of the very low-grade metamorphic zones that would permit recognition of the various transitional stages from diagenesis into metamorphism. For this reason, a separate section in this chapter has been devoted to burial metamorphism and related topics (see pp. 395-422). As to using experimental results for the interpretation of natural geologic phenomena, certain difficulties have been encountered already. Blatt (1966), for example, pointed out that altho.u,gh intergranular suturing due to pressure solution is very common in quartz-rich sandstones even of shallow burial, it has not been duplicated in the laboratory in the absence of recrystallization. It seems that extended geologic time and intergranular solutions are also required. Granulation of quartz grains, according to Blatt, was produced experimentally under uniaxial stress with piston pressures of about 400 bars. This is equivalent t o a 6,000-ft overburden, if one assumes a 1bar/l5 f t gradient. As Blatt pointed out, these results are invalid as guides t o granulation of quartz, because grains in natural deposits are not confined laterally as they are in the cylinders. Review of the literature indicates that granulation is a relatively rare compaction feature in sandstones. The information on textures in sandstones presented above demonstrates that many petrologists have concerned themselves with both the nomenclature of textures and the chemical and physical variables that control their origin. Observations on natural sand deposits, as well as on those prepared by the free-fall method under laboratory conditions, have provided data on textures and mass properties of sands that are in the initial, uncompacted state. Although the degree and style of packing and the porosity of freshly accumulated sand range widely, depending on a number of variables that must be further investigated in the future, sufficient data is available at present t o permit the sedimentologist and petrologist to define with a certain degree of reliability the characteristics of uncompacted sands of comparatively simple primary texture and composition. The data gathered from investigations on naturally and artificially compacted sands also has reached a stage where the information can be successfully used in unravelling the compaction history, preferentially as part of the whole diagenetic history.

188

K.H. WOLF AND G.V. CHILINGARIAN

POROSITY AND PERMEABILITY

The investigation of compaction of sand and sandstone, as well as any other sediment and sedimentary rock, rests partly on the study of porosity and permeability changes during the various stages of diagenesis. In this section, therefore, some fundamental data on porosity and permeability of coarser sediments, with occasional reference to clay- and silt-sized material, are presented. Although not all of this information has been directly related to compaction, a consideration of at least some selectively and preferentially chosen material on porosity and permeability is paramount in understanding and discussing compaction, and in planning future investigationsof the various compaction mechanisms and their effects. As Peck (1967;see also Pandey et al., 1974)has pointed out, there are no substances totally “non-porous” and “impermeable”, because there is no definite boundary between “diffusion”, on one hand, and “free flow”, on the other. The properties of sediments affecting diffusion and free flow are very important in the study of numerous aspects of sedimentology and petrology (e.g., diagenesis). No details on diffusion and free flow are discussed in this chapter. It is important to mention, however, that Darcy’s Law fails in extreme fluid-solids interactions, e.g., in the case of aqueous solutions of certain electrolytes flowing through some clays (Lutz and Kemper, 1959),or for fluid-solids systems characterized by Reynolds numbers greater than about 1 (Scheidegger, 1957,p. 124). Very promising seems to be a future study of the relationship between composition, texture, degree of compaction, etc., of sediments, on one hand, and the movement of fluids and ions by diffusion and free flow, on the other. In particular, the influence of the degree of compaction on porosity and permeability would be of interest as it would throw light on a number of genetic problems in the geochemistry of secondary products in sediments, including ores in detrital and carbonate rocks. As to the problems of diffusion, Pandey et al. (1974)presented fundamental theoretical concepts on pressure solution and movements of fluids. They pointed out that the migration and accumulation of solutions responsible for diagenesis and ore mineralization, as well as of oil and gas, is the result of a combination of upward and/or downward percolation, lateral movement, and diffusional processes. The reader should note that although little reference is being made to diffusion per se in this chapter, its fundamental significance is, of course, recognized. The lack of published data on the application of diffusion concepts in sedimentary petrology, is partly a reflection of the information available. As Pandey et al. mentioned, diffusion of fluid constituents is related to the configuration (e.g., size and tortuosity) of the pores within rocks. Thus, in petrologic investigations of diffusion, all

DIAGENESIS OF SANDSTONES AND COMPACTION

189

the factors discussed in the section on textures must be given due consideration. All mass properties of sandstones, including porosity and permeability, are a function of numerous variables most of which are, in turn, controlled by compactional diagenesis. Most of these variables have been treated in the different sections of this chapter. Inasmuch as porosity and permeability are among the most fundamental mass properties of sandstones, being also of considerable practical importance in the exploration for hydrocarbons, and, more recently, have been given increasing attention in ore petrology, some separate fundamental information is supplied below. Figure 3-97 (Pettijohn et al., 1972) indicates a wide variation of permeability, some in the order of 100,000 times, depending on the petrographic properties of the rock. Small-scale variations in the specific permeability values can be large in specific sandstones, especially in those that have been cemented, and usually this variability exceeds that of porosity. Tables 3-XXIII and 3-XXIV summarize the rock properties that influence permeability and flow response, wheras Table 3-XXV presents the hierarchical sequence of primary controls on permeability that can be represented also in a more complex conceptualized manner, as done by Wolf (1973a, p. 173, fig. 9), for example. The reader is referred to Pettijohn et al. (1972, pp. 93-97 and 523-533) for details on porosity and permeability as well as for examples discussed. The data on porosity and permeability is voluminous and is based on: (1) theoretical considerations; (2) experimental laboratory investigations using idealized models and natural sediments or sedimentary rocks; and (3) field studies. Some of the results are contradictory. A summary of the results of attempts to demonstrate a relationship between porosity and various fundamental textural properties is presented in Table 3-XXVI. The contradictory results, according to Griffiths (1967b, p. 230) “arise from the interdependencies between the measured properties of sediments, which, in turn, reflect the interactions between aggregates and environmental conditions in both laboratory and field observations”.

105

Specific Permeability, K Darcysl 10 1 10-1 10-2 1w3 10-4 10-5 Very fine rands silts. Clcon rand,, mixturesat sand silt bnd Unweathwed hxturssdckan

i o ~ 103 102

clay; glacial lillj&tmthed pmvcls clays; etc.

Flow charmkrirticlr

Gmd aquifers

Poor aquifers

clays

Imp.rvwus

K.H. WOLF AND G.V. CHILINGARIAN

190

TABLE 3-XXIII Permeabilities of various rock types - average values for k and K (after Davis and De West, 1966,p. 164) Type of rocks

h(mil1idarcys)

K(cm/sec)

Gravel Clean sands (good aquifers) Clayey sands, fine sands (poor aquifers) Representative values fork and K Argillaceous limestone (2%porosity) Limestone (16% porosity) Silty sandstone (12% porosity) Coarse sandstone (12% porosity) Sandstone (29% porosity) Very fine, well-sorted sand Medium, very well-sorted sand Coarse, very well-sorted sand Very well-sorted gravel Montmorillonite clay Kaolinite clay

106 - 108 103 - 106 1 -103 Wmd)

1 - 102 10-3 - 1 10-6 - 1 0 - ~ K (meinzers) * 1.8 . 10-3

* 1 meinzer = 4.72 .

cm/sec x 5.5

1.103

1.4 * 10' 2.6 1.1. 103 2.4.103 9.9.103 2.6.105 3.1 . lo6 4.3 . 107 10-2 1

-

2.50

4.74.10-2 19.90 43.60 18.00 . 10 4.6 . 103 5.8 . 104 7.88 . 105 10-4 10-2

darcys for water at 60°F.

Studies by Rittenhouse and Fraser Rittenhouse (1971a), who dealt with some of the most fundamental theoretical aspects related t o compaction of sand, has considered the amount of pore-space reduction as a result of pressure solution of quartz grains and the additional reduction of porosity caused by the chemical precipitation of the dissolved material. On considering single grains and spheres, and other idealized geometric forms, stacked according t o various packing arrangements, he found that the relative amounts of porosity loss due to solution and to cementation vary greatly, being dependent on grain shape and angularity, packing direction from which pressure was applied, and the amount of solution that had occurred. Using uniformly-sized spheres, and assuming a cubic and orthorhombic packing (as shown in Figs. 3-98a+), Rittenhouse calculated loss of pore space by solution and cementation as a result of overburden pressure (Figs. 3-98e-h). For various percentages of the radii of spheres lost by solution at points of contact, he presented (a) original and remaining porosity, (b) loss of porosity due t o closer packing, and (c) loss of porosity caused by chemical precipitation of the dissolved material as cement (Figs. 3-99a-d). The amount dissolved caused by assumed pressure solution is expressed by the percentage of the radii of the spheres lost during solution at points of contacts. For example, in fig. 3-99a, considering a loss of 28% of the

DIAGENESIS O F SANDSTONES AND COMPACTION

191

radii due t o solution, the removal of solid matter leads to a reduction in porosity of 13.3%.Inasmuch as the original porosity for this system of packing is 47.6%, and the 28% along the x-axis is equivalent to 34.3% on the y-axis, then 47.6 minus 34.3 gives the value of 13.3%lost. If the dissolved matter had been removed from the system, the only reduction in porosity would have been TABLE 3-XXIV Rock properties and flow response (after Pettijohn et al., 1972, table 11-9, p. 525) Rock property Effects on permeability and porosity Texture Grain size Sorting Packing Fabric Cement

-

permeability decreases with grain size; porosity unchanged permeability and porosity decrease as sorting becomes poorer although little data is available, tighter packing favors both lesser permeability and porosity in the absence of lamination, controls anisotropy of permeability; permeability is maximum parallel to the mean shape fabric the more cement, the less permeability and porosity

Sedimentary structures Parting lineation maximum permeability most probably parallels fabric in plane of bedding Cross-bedding scant available data suggest that horizontal permeability parallels direction of inclination and that the steeper the dip of the foreset, the weaker the horizontal vector of permeability little data, but fine grain size and more laminations combine to Ripple mark cause low permeability and, hence, ripple zones are commonly barriers to flow Grooves and flutes as judged by fabric, permeability should parallel long dimension Slump structures no data, but probably always greatly reduce horizontal permeability Biogenic structures destroy depositional fabric and bedding and, thus, drastically reduce permeability and cause minimal, if any, horizontal anisotropy of permeability; effect on porosity is unknown, but may be negligible Lithology Sandstone

Shale

thicker beds tend to be coarser grained and thus more permeable, i f cement is not a factor; if weakly cemented or uncemented, ratio of maximum to minimum permeability is perhaps less than 5 to 1; if cement controlled, ratio may reach 100 to 1 or more the prime barrier to flow that outshadows all others by far; thus it is the arrangement of sand and shale much more than permeability variation within the sand that controls flow in most reservoirs

TABLE 3-XXV Hierarchical sequence of primary controls on permeability (after Pettijohn et al., 1972, table 11-10, p. 526) Control

Remarks

Texture and fabric Defined by grain size, sorting, packing and shape orientation of framework grains Scale: 1 to a few cm3

fundamental “building blocks” that define the primary pore system; depositional fabric may be completely destroyed by burrowing organisms

Sedimentary structures Cross-bedding, ripple mark, and parting linea- directional structures consist of anisotroption are most common and nearly always have ic fabrics so that individual structures should behave as “flow packets” preferred orientation and anisotropic fabrics Scale: 1-102 m3 Bedding facies Defined by bed thickness, types and abundances of sedimentary structures and frequency of shale beds. Scale: 102-105 m3

probably the most important primary control on permeability distribution in a sandstone body; shale beds act as impermeable barriers to flow and are one of the more continuous lithologies

Composite sand bodies Superposition of one “cycle” of sand upon another, cycles commonly separated by unconformities Scale: I O ~ - I O ’ ~ m3

characteristic of many alluvial and deltaic sands; multilateral as well as multistory bodies possible

TABLE 3-XXVI Effect of fundamental properties on variation in porosity (after Rosenfeld, 1950; in: Griffiths, 1967b, table 11.4, p. 229) Rock property

Source of information* ~

theoretical analysis Coarse grain size Good size sorting Slight skewness toward finer sizes High “sphericity” and “roundness” (confused) Intermediate “roundness” Open packing Low chemical cement Low clay content

artificial mixtures

natural sediments

O+-

O+-

++

O+-

+ +

+

+

? ? ?

+ + ? ?

+

+

* 0, +, - indicate relative change in porosity for each rock property assuming the others are constant.

193

DIAGENESIS OF SANDSTONES AND COMPACTION

0

---- c u m PACKING

#--tunic PACKING, SOLUTION FROY 2 IUPPER e LOWER) CONTACTS EACH SPHERE

6 CONTACTS /SPHERE

r --cueic

PACKING, ROTATEO 45. SOLUTION FROY 4 CONTACTS OT EACH SPHERE

PACKING, ROTATLO 45. 6 CONTACTS / SPHERE

b ---CUBIC

c ---WITHORHOYBIC

OF

q --ORTHORHOYBIC PACKING, SOLUTION FROY 4 CONTACTS OF EACH SPWERE

PICKING

8 CONTACTS I SPHERE

[@ L-

_ A _ _ -

d--0RTHORHOMBIC PACKING, ROTATED 8 CONTACTS / SPHERE

WT

h --WTHORHOYBIC PACKING, ROTATE0 3 0' SOLUTION FROY 6 CONTACTS OF EACH

SPHERE

Fig. 3-98. Spheres in different packing and orientation relative to application of vertical pressure before and after solution from points of grain contacts. (After Rittenhouse, 1971a, fig. 1, p. 81; courtesy Am. Assoc. Pet. Geologists.)

194

K.H. WOLF AND G.V. CHILINGARIAN

13.376, but if all the dissolved material was precipitated in the pores, an additional loss of porosity of 9.1% (34.3 minus 25.2) would occur as calculated by Rittenhouse. Figure 3-100, on the other hand, shows that all porosity would have been lost before 30% of the spheres' radii had been dissolved. In this case, about half of the loss would be the result of solution and

47.6 X

F

F z

2

w

W

P t

!i

t > t

>

I-

a

0

P

B

a

P 0

0

IW PERCENT OF SPHERE RADIUS

LOST

ev

0

I5

SOLUTION

( b ) - U I B l C PICKING,ROTATEO 48 ' SOUITION FROM 4 CONTACTS FROM LACH SPHLRL.

( 0 )--CUBIC,VERTICAL PICKINO, SOLUTION fRw 2 COWTACTS OF EACH SPHERE.

LEGEND

LOSS 61 SOLUTlON

LOSS BY CEMENT

n

r J

50

REMAlNlNO n n E SPACE

F z W

.... , ..

X

P

W

0,

>

I-

8 B

I

0

0 PERCENT OF SPHERE RADIUS

(C

1--

OIITHORI(OMBIC PICKING; SOUITION FROM 4 CONTACTS Of EACH SPHERE.

Losr

m y SOLUTION

( d ) - - ORTHORHOMBIC PACKING; W T I O N

FROM 6 CONTACTS Of EACH SPHERE

Fig. 3-99. Original porosity and loss of pore space by solution and cementation for sphere packing and orientation presented in Fig. 3-98. (After Rittenhouse, 1971a, fig. 3, p. 83; courtesy Am. Assoc. Pet. Geologists.)

DIAGENESIS OF SANDSTONES AND COMPACTION

195

ORTHORHOMBIC, ROTATE0 30’

CUBIC,ROTATEO 45‘

POIOSITV

LOSS

DUE

TO SOLUTION

IPERCENTI

22.

.I .

OBLAlL 0 P R O L I T 1 SPHERIODS-TETRAGONAL PACKING ISPUASIIED.VERllClL,CUBlC~

20. O . O B t A l E WERMD K)*1015 SEMI A X E S - 4 C O N l A C l S 0 W L U i SPUEROIO l O l l O l 7 SEMI LXLS-4CONTACTS

I

5 W

Y

-

J

10. ’

0. (I.

a.

0

2

4

6

LORO.1TI

8 LOSS

K) DUE

I2 TO

I4

IS

SOLUTION

I8

20

22

24

ILEIISENTI

(bl

Fig. 3-100. Relationship between the loss of porosity by solution and the porosity loss resulting from the precipitation of cement. Curves in top figure are for perfect spheres, i.e., for particles with no angularity. In bottom figure, the graph for cubic-vertical packing is for oblate spheroids with perfect rounding, whereas the remaining three graphs are for prolate spheroids with angularity, the surface irregularities being expressed as minute “wedges”, “cones”,’and “pyramids” (see Fig. 3-102). (After Rittenhouse, 1971a, fig. 2, p. 82; courtesy Am. Assoc. Pet. Geologists.)

196

K.H. WOLF AND G.V. CHILINGARIAN

another half would be due to cementation. Similar calculations were made by Rittenhouse in preparing Fig. 3-99 for spheres of various packing arrangements, corresponding to those shown in Fig. 3-98.From the graphs, the cement-to-solution ratios (the relative losses by precipitation and by solution) can be determined. With greater amount of cement precipitated, the greater is the degree of pressure solution that took place. As recognized by Van Hise (1904,pp. 865-868), even the most concentrated subsurface brines can precipitate only small amounts of cement. Consequently, continued precipitation could not have come from a closed system without replenishment of solution and a steady supply of chemical elements is required. Figure 3-101 (Pettijohn et al., 1972, pp. 397-398) illustrates that with continued cementation pore spaces become progressively smaller, the permeability is steadily reduced, the flow rates decrease, and the rate of precipitation is slowed down. Adams (1964,pp. 1575-1577) indicated that the grain-size distribution of a sediment controls the initial permeability and its rate of decrease as a result of cementation. Fine-grained sands, with a lower primary permeability, undergo cementation prior to coarsegrained sediments. As indicated in Fig. 3-101,whatever flow model is proposed for the solutions causing the infilling of pore spaces by cementation, all have in common an exponential decrease in the amount of cement per

Fig. 3-101. Exponential decrease in porosity (A) and amount of cement precipitated (B), as cementation of a sandstone proceeds. The precise shapes of such curves depend on the rate of flow of solutions through the sand, the initial porosity and permeability, concentration of solutions, and the rate’of precipitation. (After Pettijohn et al., 1972, fig. 10-4, p. 398; courtesy Springer, New York.)

DIAGENESIS OF SANDSTONES AND COMPACTION

197

unit time, i.e., a long time is required for complete cementation in contrast to the time needed for partial cementation. As mentioned by Pettijohn et al. (1972, p. 398), the absolute length of time for complete cementation cannot be determined by any presently-known technique, but it has been observed that few Recent or Tertiary sandstones are completely cemented. A long period, of the order of lo8 years, seems t o be a requisite. The importance of packing in determining the cement-to-solution ratios is also shown by Fig. 3-99,a. If all four packing systems have a remaining porosity of 25%, the following ratios are applicable: cement-to-solution ratio packing system cubic-vertical 9.2113.4 cubic-rotated 4.2118.4 orthorhombic 4.2110.3 orthorhombic-rotated 4.5110.0 The calculations are based on the fact that in the first two packing systems the original porosity was 47.6% and in the other two, 39.5% (Fig. 3-100). If in all four systems only 25% porosity was left after diagenesis, then 22.6% (47.6 minus 25) and 14.5% (39.5 minus 25), respectively, were the corresponding reductions in porosity due to cementation and solution. Using the graphs in Fig. 3-994, the sum of loss of porosity by solution and that by precipitation of cement represents the diagenetic loss in pore space (e.g., for cubic-vertical packing system, as shown in the tabulation above, the total loss in pore space is equal to 9.2 + 13.4 = 22.6% and for orthorhombic packing it is equal to 4.2 + 10.3 = 14.5%). It can be seen then, as Rittenhouse pointed out, that in three of the four packing models, less than 10% cement is the result of pressure solution in cases where the quartzite has retained 25% porosity. Allen (1969, 1970) has pointed out that in theoretical considerations of problems related to textures and fabrics of natural sediments, the analogies to spherical particles of artificial sediments will give wrong results. The closest approximation can be found when oblate spheroids (or prolate spheroids) are used (Fig. 3-102), which more closely represent the spheroidal grains of the natural sands having sphericities of 0.7-0.8. Calculating the amounts dissolved by approximate methods and by using oblate spheroids of varied flatness in two arrangements, presented in Fig. 3-99,b, and a prolate spheroid in vertical arrangement, the reduction in pore space due to precipitation and solution has been approximated by Rittenhouse (1971a, p. 82) (see Fig. 3-103). These spheroid packings have the same original porosity values as those for spheres in similar arrangements. In his initial calculations, Rittenhouse only treated perfectly rounded grains; however, natural sand grains, although often well rounded as in the case of

K.H. WOLF AND G.V. CHILINGARIAN

198

Fig. 3-102. Two arrangements of oblate spheroids for which cement and solution relations were approximated. (After Rittenhouse, 1971a, fig. 4, p. 84;courtesy Am. Assoc. Pet. Geologists.)

multiple-cycle quartz particles, commonly have angularity. The “corners” and “edges” on surfaces of the grains can be visualized or expressed as minute comers or pyramids and as three-sided prisms. Results of Rittenhouse’s calculations of the cement-to-solution ratios for these types of grains 22

-

20-

I

ARE

POROSITY

LOSS

DUE

TO SOLUTION

(PERCENT)

Fig. 3-103. Relationship between maxima cement content and porosity loss due to solution for various types of sahds. (After Rittenhouse, 1971a, fig. 6, p. 85; courtesy Am. Assoc. Pet. Geologists.)

DIAGENESIS OF SANDSTONES AND COMPACTION

199

are given in Fig. 3-99,b. He pointed out that particles with rounded corners would give rise to graphs intermediate between those for spheres or spheroids and those for grains having angular corners and/or edges. Angularity decreases the cement-to-solution ratios. The original porosity increases with increasing angularity, but angular sands lose their porosity more rapidly by pressure solution at points of contact than by cementation. Rittenhouse (1971a) pointed out that pressure solution may be differential because of varying solubility from grain to grain and within individual grains. As a result, there may be penetration of the less soluble grains into the more soluble grains. In the idealized models described above, the grain and pore geometry were uniform in each case of packing; this was followed by considerations of nonsymmetrical packing and grain-size sorting and variations in other parameters. Rittenhouse (1971a, p. 86) concluded, after discussing the details, that “no type of packing and no variations in grain sphericity and roundness, sorting, or composition yield cement-to-solution ratios that are larger than found for equal-sized spheres in orthorhombic packing rotated 30”” (Fig. 3-99,a). This is in agreement with the results of Allen (1969). Thus, as pointed out by Rittenhouse (p. 87), “it appears that the ratios for this packing can be used t o indicate the maximum amount of ‘cement’ that can be derived from solution at points of grain contacts for any given amount of pore-space reduction.” If the original porosity of a rock is reduced diagenetically from 40 to 25% (a 15%reduction in porosity), the maximum amount of cement formed as a result of pressure solution (Fig. 3-104) will be 4.7%. According to Rittenhouse, his above-reviewed relationships for ideal systems can be applied to natural sandstones if certain adjustments are made. In ideal cases where the sand bodies are homogeneous and where all the dissolved material was precipitated in adjacent pores, “the cement-to-solution ratios for rotated orthorhombic packing appear to apply best to wellrounded, well-sorted sands that had original porosities of about 38%. For more angular sands, for example, for sands having higher or lower porosities, or for those having better or poorer sorting, these cement-to-solution ratios appear to be high.” Rittenhouse estimated maximum cement-to-solution ratios, as given in Fig. 3-103, for the following: (1)very poorly-sorted sand (porosity about 28%); (2) extremely well-sorted sands (porosity about 42.5%); and (3) well-sorted, very angular sands (porosity about 43%). The maximum values of cement-to-solution ratios for sands with intermediate values in porosity, sorting, and angularity would fall between the two curves plotted in Fig. 3-103. This “best estimate’’ curve of Rittenhouse is the same graph as the one labelled “wedges” in Fig. 3-99,b. Taylor (1950) already had observed the behavior of volcanic and other ductile grains under pressure (e.g., shale, anhydrite, and salt) and that their

K.H. WOLF AND G.V. CHILINGARIAN

200

nI2

-

10-

I-

well sorted rond.

6-

Very poorly sorted (2) Ettremely well rorlcd,of

(1 4-

(3) Very onqulor 29

0

2

4

6

8

10

ORIGINAL

12

I4

MINUS

16

18

20

22

24

26

28

30

32

34

P R E S E N T POROSITY [ P E R C E N T )

Fig. 3-104. Relationship between maxima cement content and difference between original and present porosity for various sands. (After Rittenhouse, 1971a, fig. 7,p. 86; courtesy Am. Assoc. Pet. Geologists.)

deformation increased with depth of burial. According to Rittenhouse (1971b, p. 92), the available experimental data suggest that sandstones containing ductile grains may undergo more intense mechanical compaction than those devoid of ductile grains and rich in quartz and/or feldspar and might lose porosity and permeability more rapidly. An example of this type of sands would be graywackes (Folk’s phyllarenites, i.e., a rock rich in phyllite rock fragments) containing relatively “soft” sedimentary and/or metamorphic rock fragments. Sawabini et al. (1974) also showed that compressibility increases with increasing feldspar/quartz ratio. It has been suggested by numerous investigators that postdepositional deformation due to compaction of the ductile fragments may give rise to a “matrix” (see Dickinson, 1970, for example). Rittenhouse (1971b) has offered an ideal model of a sand composed of nonductile and ductile spherical grains. The results of compactional deformation of the “soft” grains leading to various degrees of pore space and thickness reductions are presented in Figs. 3-105 and 3-106. Fraser (1935) discussed packing and its relationship to porosity and permeability in great detail (see also Graton and Fraser, 1935).When a number of spheres of constant size are uniformly packed, the remaining inter-sphere openings can be occupied by spheres that are of a particular smaller size. This can be repeated until the unit voids are filled with successively smaller

201

DIAGENESIS OF SANDSTONES AND COMPACTION

kEDUCTlON I N PORE SPACE,

Fig. 3-105. Relation between content of ductile grains and reduction in pore space caused by compaction. (After Rittenhouse, 1971b, fig. 2, p. 92; courtesy Am. Assoc. Pet. Geologists.)

r 45 40

a-' VI

3

35

30

4

f

25

W

10

5 0

0

5

10

I5

20

25

30

35

40

REDUCTION I N THICKNESS{COMPACTlON), o/'

Fig. 3-106. Relation between content of ductile grains and reduction in thickness caused by compaction. (After Rittenhouse, 1971b, fig. 3, p. 95; courtesy Am. Assoc. Pet. Geologists.)

202

K.H. WOLF AND G.V. CHILINGARIAN

20

24

28

32

36

40

POROSITY OF BINARY MIXTURE, YO

Fig. 3-107. Relationship between the matrix grain size (expressed as a fraction of diameter of large spheres) and the porosity of binary mixture. (After Fraser, 1935, fig. 1; by permission of The Univ. of Chicago Press, copyright @ Univ. of Chicago.)

and smaller spheres. Fraser stated (p. 921) that if the voids in an assemblage of spheres are filled successively with smaller spheres, the resulting porosity values for the combination (mixture of spheres) do not fall on a smooth curve. Instead, the porosity decreases by sudden steps, as shown in Fig. 3-107. As to mixtures of spheres of two sizes, Fraser stated that as long as the proportion of the smaller spheres is sufficient to keep the larger spheres separated from each other, the smaller spheres will dominate the general fabric or structure of the assemblage. The small spheres together with their own interparticle voids represent the “matrix”, whereas the larger spheres can be considered to be “foreign” particles floating in the matrix and disturbing the assemblage. Fraser noticed a fairly uniform decrease in porosity with increasing proportion of the large spheres (for an explanation, see p. 918 in his publication). According to Fraser, when the proportion of large spheres is increased beyond a certain limit, two alternative situations occur, depending on the relative diameters of the assemblages of spheres: (1)When the proportion of large spheres is just sufficient for them to be in contact with each other and to be self-supporting without requiring support from the small spheres, and when the smaller spheres have a diameter less than that corresponding to the “critical ratio* of occupation”, then the large spheres control the fabric or structure of the assemblage. The latter also holds true with further increase in the percentage of large spheres. At the particular point just mentioned, when the control changes from the smaller

* Ratio of the diameter of a small sphere, which can just pass through the pore throat between larger spheres into the interstitial void, to the diameter of the larger sphere.

DIAGENESIS OF SANDSTONES AND COMPACTION

203

to the larger spheres, the former just fill the voids between the latter ones without causing any distortion of the packing of the assemblage of the large spheres. Beyond this turning-point, as the proportion of large spheres increases and the smaller ones no longer can fill all the openings between the large particles, additional voids remain open. Consequently, beyond this turning-point, an increase in number of large spheres results in an increase in porosity. (2) In distinction from case 1, when the proportion of the large spheres increases beyond the limit of domination by the small particles, and when the diameters of the small spheres exceed the “critical ratio of occupation”, then and thereafter the two sets of spheres mutually interfere with each other. Neither set controls the fabric or structure of the assemblage until the content of large spheres reaches a total of 100%. Mutual interference raises the porosity; but, on the other hand, the porosity still decreases as the proportion of the large spheres increases. Certain of these relationships are shown in Fig. 3-108: Curve 1 represents porosities when the diameter of the smaller of the two series, d , is equal to 0.433 D,where D is the diameter of the larger spheres, and curve 2 depicts porosities when d is equal to 0.158 D. Curve 1 (“total porosity”) demonstrates a slower decrease in porosity with an increase in the proportion of large spheres than does curve 2, because the difference in the diameter of the spheres is greater in the latter case. When the content ratio of large to small spheres is 3/1 (= 72-7576 large spheres to 25-28% small spheres by volume), the large particles commence to control the fabric or structure of the assemblage and the porosity consequently abruptly increases (dashed lines). This porosity increase is particularly marked when the spheres differ widely in diameter. The two lines converge at the same porosity (39%) when the assemblage is composed of large spheres only (Fig. 3-108). Figure 3-108 also indicates that an increase in the proportion of large spheres results in increasing disturbance and looseness of packing of the small spheres which constitute the “matrix”. The two upper curves (“porosity in matrix”) present the proportion of remaining openings between the large spheres after their interparticle spaces have been occupied by the smaller spheres. As the number of large spheres increases, the porosity of the matrix also increases, because of growing disturbance of the fabric or structure and the loosening of the packing among the “matrix” as a result of increasing number and closeness to each other of the large spheres. The porosity of the assemblage of spheres changes with change in the ratio of the two diameters (Fig. 3-log), if the proportions of the two populations of spheres are constant. In Fraser’s experiments, the volumes of spheres of two different sizes were always equal; the diameter ratio of large to small spheres ranged from 1 : 1 (i.e., only one size of spheres) to 19 : 1. According to the data obtained by Fraser, the decrease in porosity is most marked in

K.H. WOLF AND G.V. CHILINGARIAN

204 LARGE SPHERES, %

t

t 0

d

4

so SMALL SPHERES, %

100

Y

POROSITY. %

Fig. 3-108.Relationship between the porosity and the relative proportions of large and small spheres. 1 and 3 = diameter ratio of large to small spheres is equal to 1 : 0.433;2 and 4 = diameter ratio is equal to 1 : 0.158. Curves 1 and 2 indicate porosity of matrix, whereas curves 3 and 4 show total porosity. (After Fraser, 1935, fig. 2;courtesy J. Geol.) Fig. 3-109.Relationship between porosity and the ratio of large to small sphere diameters for mixtures containing 50% of each size. (After Fraser, 1935,fig. 7; by permission of The Univ. of Chicago Press, copyright @ 1935 University of Chicago.)

the range of size ratios from approximately 2 : 1 to 6 : 1. Beyond a ratio of about 8 : 1, the growing difference in size of spheres of the two populations reduces the porosity only slowly. The effect of angularity on porosity was also considered by Fraser. He used a series of carefully sized materials, ranging in shape from spheres to flat plates, and maintained other parameters as constant as reasonably possible. The materials tested included: (1) the lead and sulphur shot (perfectly spherical sand-sized grains); (2) the marine and beach sand (partly rounded grains); (3) the crushed material (angular); and (4) the mica (flat, plate-like fragments) (Table 3-XXVII). The porosity was measured before and after the material was compacted by jarring, first when the constituents were dry and then when they were saturated with water. As shown m the Table, the angularity markedly affects the packing with a consequent increase in porosity. Assuming that the influence of grain shape on packing results in the variation of porosity, Fraser’s data showed a range of 8.73%(43.51-34.78) units (excluding mica) in porosity of compacted dry materials, i.e., from 34.78 to 43.51%. Of this range, 6.25 units are due to increase in porosity, because of irregular angularity, whereas 2.48 units are owing to porosity increase as a result of the ptesence of flattened particles (Table 3-XXVII). Moderately well-rounded sands exhibited only minor variations in porosity

DIAGENESIS OF SANDSTONES AND COMPACTION

20 5

TABLE 3-XXVII Effects of grain shape on porosity (after Fraser, 1935, p. 936) Material

Specific gravity

Porosity (%) type of packing dry loose*

~

Lead shot Sulphur shot Standard sand (marine) Beach sand Dune sand Crushed calcite Crushed quartz Crushed halite Crushed mica

11.21 2.024 2.681 2.658 2.681 2.665 2.650 2.180 2.837

wet compacted**

~~

40.06 43.38 38.52 41.17 41.17 50.50 48.13 52.05 93.53

loose

compacted

42.40 44.14 42.96 46.55 44.93 54.50 53.88

38.89 38.24 35.04 38.46 39.34 42.74 43.96

~~

37.18 37.35 34.78 36.55 37.60 40.76 41.20 43.51 86.62

~

~~

-

-

92.38

87.28

* and ** = before and after compaction by jarring. as a consequence of limited variation in the degree of rounding. The angularity led t o “bridging” and loose original packing of the grains. Independent of the grain shape, the wet constituents packed more loosely than the dry materials (Table 3-XXVII). This has also been observed in natural sand accumulations. The effect of angularity is more pronounced in the case of wet packs, as demonstrated by the differences in porosity of dry and wet packs for rounded versus angular materials (both loose and compacted). The accumulations of flat and needle-like components exhibit the greatest porosity (e.g., mica has over 90% porosity), which is also supported by observations on natural sediments. These high porosities cannot be reduced below 80%, even by prolonged shaking, as done by Fraser. An increase in pressure, causing mechanical compaction, was required to diminish the porosity of the wet mica accumulation to 67.4%. It should be pointed out, however, that the compressibility of sands increases with increasing mica and clay content (Sawabini et al., 1974). Fraser concluded that angularity usually increases porosity. He observed a decrease in porosity caused by “angularity” only in cases where the grains were mildly and uniformly disk-shaped. The control on permeability by the unformity of grain size, i.e., sorting, is very great. Just ‘as porosity is less for a population of mixed sizes, the permeability is also reduced within certain limits when particles of different

K.H. WOLF AND G.V. CHILINGARIAN

206

sizes are added. As to Fraser’s (1935) two-component populations, he remarked that a layer of openings, considered larger than those existing within the population of small spheres, exists around each large sphere. The small spheres adjacent to large ones can touch them only at one point, and the distance between these contact points is controlled by the size of the small spheres of the “matrix”. The relatively large voids around the large spheres connect freely in all directions, giving rise to a very permeable area or zone. As a result of disturbance of the packing of the “matrix” by the presence of a large “foreign” sphere, an area of looser packing extends for a distance equal to several large-sphere diameters away from these large particles. This also increases the porosity and permeability. Both of the influences mentioned above oppose the decrease in permeability caused by the presence of impermeable large spheres. The relative values of these opposing controls are given in Fig. 3-110 for a number of cases investigated by Fraser. The upper curve was obtained by adding larger spheres to a population of smaller ones with a diameter of 2.3 times smaller than that of the large spheres. With increasing proportion of the large spheres, initially the permeability increases slowly and then more rapidly. The reason for this behavior lies in the fact that the larger spheres are more effective in increasing the channelways by disturbing the packing of smaller spheres than in decreasing the permeability by blocking previously existing channels. It is rather interesting to observe that whereas the total pore space that remains after each addition of large spheres is reduced, the remaining pores form larger and more effective chanLARGE SPHERES, Yo

0;

‘

’

’

’

’

50

’

‘

SMALL SPHERES, %

’

‘

’

100

Fig. 3-11 0 . Relationship between the coefficient of permeability and the relative proportion of large and small spheres. 1 = ratio of the diameter of large sphere to that of the small one is equal to 3.61; 2 = large/small sphere diameter ratio is 6.28; 3 = large/small sphere diameter ratio is 2.30.. (After Fraser, 1935, fig. 11; by permission of The Univ. of Chicago Press, copyright 01935 Univ. of Chicago.)

DIAGENESIS OF SANDSTONES AND COMPACTION

20 7

nels. This creates the anomaly of increasing permeability accompanied by a decreasing porosity . As the lower curve in Fig. 3-110 indicates, the adding of large spheres with a relative diameter of 6.23 (i.e., diameter of large spheres is 6.23 times larger than that of small spheres) t o a population of smaller spheres initially causes a lowering of the permeability, which continues until the ratio of large to small spheres is about 50 : 50 by volume percent. Thereafter, the permeability increases until the content of the smaller spheres reaches about 31.5%by volume, when the permeability of the mixture is equal to that of the small spheres alone. From this point on, the permeability is greater. On the other hand, the porosity of the two-component system decreases until the ratio of the small to large spheres is 25 : 75. After that, the large spheres are in contact with each other and the spaces between them are completely filled with the small spheres. It is obvious then, that in mixtures containing from 25 to 50% of small spheres, the total pore space decreases whereas the permeability increases with increasing content of larger spheres. When the content of small spheres falls below 25%, the value of the porosity loses significance, because the small spheres are present in insignificant numbers to fill all the interstices between the large spheres. The central curve in Fig. 3-110 was obtained for a two-component system containing spheres, diameters of which were in the ratio of 3.6 : 1. Fraser (1935, pp. 941-946) discussed compaction of sands to some extent. He mentioned that it has been stated that sands show no shrinkage on drying and that this appears to be incorrect. As mentioned earlier, wet sand packs less tightly than dry sand by 1%or more percentage units of the total porosity (Table 3-XXVII). Experiments showed that wet sand had a porosity of 37.92%, whereas the same sand when dried had a porosity of 36.95%,and these results were reproducible. The difference may be due to the removal of a film of water that is present around each sand grain. Fraser discussed the degree of compaction of beach sands and stated that any accumulation of detrital components has a fairly definite range of porosity reflecting various degrees in the perfection of packing of the grains. He found that a sample of a moderately well-sized and well-rounded beach sand in the dry state exhibits a range in porosity from 37.8% to 46.576, depending on whether the fabric is that of a loose or tight packing. When wet, approximately the same range of porosity is present. The above change in porosity represents a change of 14%in the total volume of the sand. More research data is needed for better understanding of the influence of the numerous environmental parameters on the original porosity and the subsequent diagenetic compactional history as a result of increasing overburden. In one of his experiments, Fraser used actud beach sand that had an average natural porosity of 40.56%prior to sampling. Mere settling of the sand in the laboratory into a

K.H. WOLF AND G.V. CHILINGARIAN

208

container produced a porosity of 46.0% when wet and 46.3% when dry. Tapping produced settling or “compaction” until after 6 minutes the constant values of 38.26%and 37.82%in the wet and dry states, respectively, were obtained. These observations indicate that the naturally-accumulated sand on the beach with an average porosity of 40.56%had already undergone sufficient rearrangement to have lost most of the bridging effect. It has assumed such a fabric that further compaction while wet could take place only slowly, although considerable additional consolidation and loss of volume on drying still existed. Fraser also found that dry sand may be compacted much more quickly, especially during the early stages, than the wet sand (see table 6 in Fraser, 1935). Fraser further observed that when bridging of the sand grains is not present, no significant compaction can take place until the pressure applied to the system exceeds the crushing strength of the minerals. In the coarse sediments, the grain-size distribution and the matrix content are very important in controlling compaction. The degree of “compaction” mentioned above is, of course, related to the experimental conditions used by Fraser in attempting to cause the rearrangement of the fabric by tapping; therefore, the results cannot be easily extrapolated to natural or laboratory conditions where the sand is exposed to higher pressures and temperatures and to intrastratal fluids of varying composition. Experiments such as those done by Fraser should be performed at high temperatures and confining pressures. In addition, the influence of sorting, orientation, mineral composition, and grain shapes on compressibility, porosity, and permeability should be investigated in greater detail. The permeabilities of gravel packs having a porosity of about 35% are presented in Table 3-XXVIII.According to Hill (1941, p. 138), flow turbulence in part determines the permeability of gravels. Porosity also plays an important role, e.g., an increase in porosity from 35 to 40%results in an increase in permeability by as much as 60%. TABLE 3-XXVIII Permeability ranges of gravel packs (Q,% 35%) for gravels of different sizes (after Hill, 1941,table 3, p. 138) Gravel size (mesh)

Range in permeability (darcys)

Average permeability (darcys)

3- 4 4- 6 6-8 a-1 0 10-14

72 00-9000 3400-4000 1700-21 00 1000-1’300 700- 900

8100 3700 1900 1150 800

DIAGENESIS OF SANDSTONES AND COMPACTION

209

TABLE 3-XXIX Permeabilities o f various gravel-sand mixtures (after Hill, 1941, table 4, p. 129) ~

~~~

~~

Gravel size (mesh)

Gravel/sand size ratio

Average gravel permeability (darcys)

(Gravel + sand) permeability (darcys)

3- 4 4- 6 6- 8 8-1 0 10-14

15.0 10.6 7.5 5.3 3.7

8100 3700 1900 1150 800

40- 80 180-220 230-300 250-300 200-250

(avg. 60) (avg. 200) (avg. 265) (avg. 275) (avg. 225)

Hill (1941, p. 139) determined the permeability of these gravels after bridging them with a sand having the following properties: (1) logarithmic mean = 1.863 mm; (2) geometric mean = 0.275 mm; (3) standard deviation = 0.674; (4) skewness = 0.268; (5) kurtosis = 0.100; (6) ten percentile = 0.445 mm; (7) permeability = 6.6 darcys; (8) porosity = 35%.Sand was introduced into the gravel in a vertical-flow tube by flowing sand-water mixture. The results are presented in Table 3-XXIX. The permeability of the gravel-sand mixture was highest (average k = 275 darcys) in the case of gravel/sand size ratio of 5.3, whereas the lowest permeability (average h = 60 darcys) was present when the gravel/sand size ratio was 15. In the latter case there is an intensive migration of sand into the gravel caused by poor bridging action. Landreth (1969, p. 4) also presented permeabilities of packed gravels of certain size ranges (Table 3-XXX). Gaither (1953) considered the effects of sorting on porosity which can be quite complicated. Commonly, the porosity diminishes as the grains deviate from uniform size distribution, because: (a) the finer grains fill the voids TABLE 3-XXX Relationship between the particle size and permeability of gravels artificially packed in oil wells in order t o preclude sand production (after Landreth, 1969, p. 4) Gravel size range (inches)

U.S. Series Number (mesh)

Permeability (darcys)

0.023-0.032 0.032-0.046 0.046-0.065 0.065-0.093 0.093-0.1 31 0.131-0.18 5 0.185-0.2 63

20-30 16-20 12-16 8-1 2 6- 8 4- 6 3- 4

800 1100 1500 2100 2700 4000 6500

K.H. WOLF AND G.V. CHILINGARIAN

210

between the larger grains to form what is generally known as “matrixyy,and (b) the coarsest grains reduce the porosity by occupying a volume that would otherwise be occupied by the finer, porous material. These findings are in agreement with the observations made by Fraser (1935),as discussed above. Using King’s (1898)experimental data on porosities of sands, Gaither presented porosity curves for well-sorted sand and mixtures of sands (Fig. 3-111).Using a mathematical approach, Von Engelhardt (1960)presented an idealized (theoretically determined) curve. Examination of Fig. 3-111shows that (a) well-rounded, well-sorted sands have porosities ranging from about 34% for coarse sands to 38% for fine sands (curve 1); (b) decreasing the sorting by mixing different proportions of two different-sized sands results in a porosity decrease (curves 2 and 3 which are below curve 1); and (c) the lowest porosity of around 25% corresponds to a mixture of 60--70% sand finer than 0.096 mm in diameter and 30-40% of a well-sorted sand with an average diameter of 0.483 mm (curve 3). The value of 25% is approximately 12% below the average for a well-sorted, well-rounded, medium-sized sand. Morrow et al. (1969)presented information on the prediction of porosity and permeability from grain-size data. They pointed out that on the basis of a systematic study of unconsolidated material, Fumas (1929)showed that: (1) the porosity of a two-component sediment depends on both aggregate

0%

I

COARSE

100%

Fig. 3-111. Variations in the porosity with varying proportions of fine-grained and coarse-grained sands. 1 = change of porosity with average grain size for well-sorted sands (“simple” sands); 2 = change of porosity in mixtures of sands with average diameters of 0.611 mm and 0.152 mm; 3 = change of porosity in sand mixtures in which average diameter of coarse sands is 0.483 mm and maximum diameter of fine sand is 0.096 mm. (After King, 1898; in Gaither, 1953, fig. 2, p. 185; courtesy J. Sed. Petrol.)

DIAGENESIS OF SANDSTONES AND COMPACTION

211

40

8

*t

!

30

20

PERCENTAGE O F COARSE-GRAINED FRACTION

Fig. 3-112.Relationship between the porosity and the relative proportions of fine and coarse fractions in two-component sand aggregates for various size ratios. Percentages are by voIume. (After Furnas, 1929;in Morrow et al., 1969, fig. 1,p. 312;courtesy J. Sed. Petrol. )

I

0

0 c

10

-

*L

(L

I

.lo

x)

40

60

80

1 1

100

PERCENTAGE O F COARSE- GRAINED FRACTION

Fig. 3-113.Relationship between the ratio of (permeability of the mixture)/(permeability of the aggregate of small components) and the relative proportions of the fine and coarse fractions for two-component sand aggregates. Radius of coarse grains = 0.1 cm; porosity ofeither fine fraction or coarse fraction = 40%. Numbers on the curves designate the ratio of (permeability of the coarse fraction)/(permeability of the fine fraction). (After Furnas, 1929;in Morrow et al., 1969,fig. 2,p. 313;courtesy J. Sed. Petrol. )

212

K.H. WOLF AND G.V. CHILINGARIAN

composition and the size ratio of the particles (which has also been mentioned by Gaither, 1953, as discussed above), but is independent of particle size (Fig. 3-112); (2) the permeability of a two-component aggregate depends on relative proportions of fine-grained and coarse-grained fractions and their respective permeabilities (Fig. 3-113). In Figs. 3-112 and 3-113, the percentage of the coarse grains in an aggregate is plotted on the abscissa and ranges from 0 to 100%. At both extremes, of course, the system becomes a one-component sediment, composed either of fine or coarse grains, so that the lines converge to meet. The smalr t o large size ratio in Fig. 3-112 ranges from 0.5 to 0.05. The ratio of 1.0 would indicate no difference in size so that the curve would be a straight line parallel to the abscissa. With an increase in difference in grain diameter, greater volume of the intergranular space between the large grains can be filled by the smaller particles, giving rise t o a distinct decrease in porosity. The maximum decrease takes place when the two-component aggregate is composed of approximately 70% coarse material and 30% fine-grained components, for a size ratio of 0.2, 0.1 and 0.05. Morrow et al. (1969) determined the interrelationship among the grain-size distribution of unconsolidated sediments, porosity, and permeability. The mathematical formula that would describe the grain-size distribution best, according to them, is the Rosin-Rammler equation:

(3-3) where Y = cummulative weight percent of under-sized material, z = particle size, N = a measure of the narrowness of the particle size distribution, and D = a measure of the mean particle size. They also found that two other measurements, in addition to N and D , were useful in their correlations, i.e., fines = actual weight percent of material that passed a 325-mesh (44-micron) screen, and RR fines = weight percent of material less than 44 microns in size given by the fit to the Rosin-Rammler equation (eq. 3-3). These two measures enabled determination of the fine-end skewness. The findings of Morrow et al. can be summarized as follows: (1)The best correlation between the grain-size distribution and porosity was given by a plot of porosity as shown in Fig. 3-114. (2) Porosity versus log[lOON(RR fine~)l/~/fines], tends to increase with increasing closeness of the particle size distribution (i.e., N). (3) Porosity is independent of absolute particle size: there was a complete scatter on a plot of porosity versus D . (4) There is a reduction in porosity with increasing fractional weight of fines, which fill the interparticle spaces of a packing matrix formed by larger particles. ( 5 ) Porosity depends on both grain-size distribution and packing geometry. (6) The good correlation between the grain-size distribution and porosity suggests that the mode of packing is reasonably consistent for a given size distribution. (7) The

DIAGENESIS O F SANDSTONES AND COMPACTION

31

213

I

t

( 100xNx(RRFINES)”3 FINES

)

Fig. 3-114. Relationship between porosity and Rosin-Rammler size-distribution parameters. Root mean-square deviation = 1.65%; overlapping points = 7. Fines = actual weight percent of material that passes a 325-mesh (44 p ) screen; R R Fines = weight percent of material less than 44 p in size given by the fit to the Rosin-Rammler equation. (After Morrow et al., 1969, fig. 6, p. 317; courtesy J. Sed. Petrol.)

0

E

h

I

f

t

c

i I

W h

Y

4 LOG

(NrD),

microns

Fig. 3-115. Relationship between permeability and Rosin-Rammler size-distribution parameters. Root mean-square deviation = 0.32; overlapping points = 12. (After Morrow et al., 1969, fig. 7, p. 318; courtesy J. Sed. Petrol.)

K.H. WOLF AND G.V. CHILINGARIAN

214

LOG [ N X D X ( R R F I N E S ) ” ~ ] , ~ ~ C ~ O ~ ~

Fig, 3-116.Relationship between the permeability and the Rosin-Rammler size-distribution parameters. Root mean-square deviation = 0.345;overlapping points = 18. (After Morrow et al., 1969,fig. 8,p. 319;courtesy J. Sed. Petrol.)

MEDIAN DIAMETER,

/4

MEDIAN DIAMETER. mrn

Fig. 3-117A.Median diameters and porosity of recent North Sea sediments of Wilhelmshaven. (After Fuchtbauer and Reineck, in: Von Engelhardt, 1960, fig. 7; courtesy Springer, Berlin.) Fig. 3-117B.Median diameter and porosity of recent shelf sediments of the Californian coast (San Diego County). (After Hamilton and Menard, in: Von Engelhardt, 1960,fig. 8; courtesy Springer, Berlin.)

DIAGENESIS OF SANDSTONES AND COMPACTION

21 5

permeability of a sand increases with increasing mean particle diameter and with increasing porosity. Relationships between the permeability and the Rosin-Rammler size-distribution parameters are given in Figs. 3-115 and 3-116. ( 8 ) The porosity varies directly with the closeness of the size distribution. (9) The relative amount of fine particles affects the permeability through interstitial blockage (see also Fig. 3-113 of Furnas, 1929). Von Engelhardt (1960) observed that the primary porosity of freshlydeposited sands is dependent on the grain size down to the range of fine sand, as shown in Figs. 3-117 and 3-118. In one case, the fractional porosity between 0.40 and 0.44 is independent of grain size in the range of 120 to 240 p, whereas in the second example the fractional porosity between 0.38 and 0.45 is independent of grain size between 200 and 700 p. This indicates that the results of experiments with spheres apply only to coarser sands and are less applicable as the grain size decreases; this cannot be explained on the basis of purely geometric considerations. The examples of high-porosity sandstones obtained from the subsurface are presented in Tables 3-XXXI and 3-XXXII. The data indicates that inspite of overburden pressures, sandstones can retain a considerable amount of pore space, unless the pores have been filled as a result of chemical precipitation. Uncemented, very friable sediments have been found even at relatively great depth (approximately 4000 m), and it is in these cases where one can demonstrate that mechanical compression alone can reduce the pore space. Table 3-XXXII and Figs. 3-118 and 3-119 show examples of sands and sandstones from oil fields which have undergone mainly the mechanical

h OD6 @X12 .I3 .B .x)40 .M) I.Omm

Fig. 3-118. Grain-size distribution, porosity ($) and permeability (k in darcys) parallel to the bedding of unconsolidated to slightly consolidated sandstones of Valendis (Bentheimer Sandstone), Germany. From drillholes of the Ruhlermoor Oil Field near Meppen/Ems, Germany. (Laboratory of the Gewerkschaft Elwerath.) A = from depth of 782-788 m, 4 = 0.296, and k = 11.0 d; B = from depth of 782-788 m, $ = 0.297, and k = 8.70 d; C = from depth of 759’800 m, 4 = 0.325, and k = 2.10 d. (After Von Engelhardt, 1960, fig. 9; courtesy Springer, Berlin.)

K.H. WOLF AND G.V. CHILINGARIAN

216

TABLE 3-XXXI Examples of deeply-buried sandstones with relatively high porosities (after Von Engelhardt, 1960, table 3) Locality

Formation

Depth (m)

Porosity fraction

Weber, Pennsylvanian Cockfield, Eocene Miocene Tennsleep, Pennsylvanian Eocene Eocene Frio, Oligocene Oil Creek, Ordovician Pliocene Bromide, Ordovician

1860 2100 2160 2280 2320 2340 2740 3260 4300 4600

0.176 0.298 0.280 0.195 0.270 0.315 0.282 0.067 0.200 0.050

Lower Pliocene Lower Pliocene Lower Pliocene Lower Pliocene Lower Pliocene Upper Miocene

1555 1575 1695 1930 1960 2530

0.360 0.232 0.317 0.280 0.273 0.300

U.S.A. (1 ) Rangely, Colorado (2) Katy, Texas (3) University Field, Louisiana (4) Big Medicine Bow, Wyoming (51 Davis Lens, Texas (6) Liberty Co., Texas (7) Fishers Reef, Texas (8) Lindsay, Oklahoma (9) Fillmore, California (10) Carter Knox Field, Oklahoma

Italy (11) Cortemaggiore near Piacenza (12) Cortemaggiore near Piacenza (13) Cortemaggiore near Piacenza (14)Cortemaggiore near Piacenza (15) Cortemaggiore near Piacenza (16) Budrio East, near Bologna

NO. 1 , 2 , 6, 7, 8 after Winsauer, Shearin, Masson, Williams (1952); No. 3, 5 after Levorsen (1956);No. 4 after Waldschmidt (1941);No. 9 after Henriksen (1958); No. 10 after Stearns (1957); No. 11-16 after AGIP (1959). TABLE 3-XXXII Unconsolidated to slighlty consolidated sandstones from German oil fields (after Von Engelhardt, 1960, table 4) Locality Eldingen near Celle Eldingen near Celle Scheerhorn near Nordhorn Scheerhorn near Nordhorn Riihlermoor near Meppen Riihlermoor near Meppen

~

Formation

Depth (m)

Porosity

Lias (Y Lias (Y Valendis Valendis Valendis Valendis

1483 1463 1104 1120 842 853

28 29 23 27 30 33

* median diameter; ** sorting coefficient.

(%I

d50*

d75/d25

so**

1.60 1.20 1.49 1.60 1.67 1.59

1.12 1.11 1.04 1.23 1.01 1.10

(mm) 0.133 0.105 0.340 0.113 0.270 0.138

DIAGENESIS OF SANDSTONES AND COMPACTION

217

Fig. 3-119. Grain-size distribution, porosity (@) and permeability (k in darcys) parallel to bedding of slightly consolidated sandstones of the Lias (11. From drillholes in the Eldingen Oil Field near Celle, Germany. (Laboratory of the Gewerkschaft Elwerath.) A = from depth of 1490 m, # = 0.31, and k = 0.950 d; B = from depth of 1483 m, q5 = 0.28, and k = 0.420 d. (After Von Engelhardt, 1960, fig. 10; courtesy Springer, Berlin.)

compaction and only little chemical cementation. The sands from Ruhlermoor (Germany) are very loose, those from Scheerhorn (Germany) only weakly cemented, and those from the Eldingen oil field (Germany) are slightly more consolidated. These sands show good sorting and are of marine origin. The median diameter lies between 0.1 and 0.34 mm and the distribution curve is symmetrical. According to practical and theoretical considerations, the true primary porosity is about 40%. Thus, without chemical infilling of pores, the porosity was reduced from 40% to 23-33%, i.e., reduction of about 20-40%. Inasmuch as the plastic deformation and breakage of quartz is not present, the reduction in pore space is due to the rearrangement of the grains, but the movement was only minor. Primary textures are either modified or destroyed. In sands in which the primary orientation of grains was still observable after pore-space reduction, rotation along the long axes must have constituted the movements during compaction. According to Von Engelhardt (1960),if sand deposits of different sizes and shapes are undergoing compaction under the same overburden pressure, the mechanical reduction of the primary pore spaces will be different: the fine sands will experience less reduction in porosity than the coarser-grained ones, as the number of contacts per unit volume in fine-grained sands is larger and the resistance to compression is greater. The shape of the grains also plays a role in controlling degree of compaction. The sands composed of grains having irregularly-shaped surfaces maintain a higher porosity in contrast to sands having smooth, spherical grains, as the former have greater

K.H. WOLF AND G.V. CHILINGARIAN

218 30

I

I

I

I

I

I

I

I

-

28

2 24v)

2

22 20

0 00

-

100

180

280

I

#

340

1

1

420

500

MEDIAN DIAMETER,/

Fig. 3-120. Porosity of different beds of the Bentheimer Sandstone (Valendis) from a drillhole of the Scheerhorn Oil Field (Lingen, Germany). The porosity is controlled by the median diameter as shown by 190 measurements. (After Von Engelhardt, 1960, fig. 13; courtesy Springer, Berlin.)

numbers of long and concavo-convex contacts. An example illustrating the influence of grain size in a case where sands of different grain sizes are present is given in Fig. 3-120. A 25 m thick sandstone unit is coarser at the top and becomes finer towards the bottom. The distribution curve, e.g., sorting, is the same, so that only grain size varies. Figure 3-120 shows that the porosity of the fine sands is on the average greater than that of the coarse sands. Pettijohn et al. (1972, pp. 93-97) mentioned that the permeability of sandstones, especially from the same lithologic bed, often show a lognormal statistical distribution, whereas porosity is characteristically normally distributed. Variations in permeability are much greater than that of porosity as shown by Fiichtbauer (1967a) in his log permeability-versus-porosity plots. The same plots reveal an obvious correlation between effective porosity and permeability. The Kozeny-Carman equation shows that permeability is inversely proportional to the second power of specific surface (see Langnes et al., 1972). As demonstrated in Fig. 3-121, as grain size decreases, specific surface of the sediment increases and the resistance to flow increases. The permeability of fine-grained sand (or silt) is smaller than that of a coarsegrained one, with the effective porosity of both sands being identical. Von Engelhardt (1960) pointed out that the nature of sediments is such that permeability is never isotropic. The analyses of fluid movements in the pore spaces must take into consideration the fact that permeability changes with changes in the direction of fluid movement. Most important are the differences between the permeability parallel and normal to the bedding. In all examined cases, the permeability of sandstones normal to the bedding is less, as shown in Fig. 3-1122. Two important questions arise here: (1)What are the changes of permeability parallel and across the bedding during me-

DIAGENESIS OF SANDSTONES AND COMPACTION

219

I'

0

.o I

LOG GRAIN SlZE,h

..a"

' I

' 0

I

100

10

I"

I

Kx)

loo0

_

I IQoOord

kll

Fig. 3-121.Relationship between the permeability ( k ) and grain size ( d ) for the Bentheimer Sandstone of the Scheerhorn Oil Field. The regression equation is log k = -2.1007 + 2.221 loglod, where k is permeability in millidarcys and d is grain size in millimeters. Scatter diagram is based on random selection of data from fig. 49 of Von Engelhardt (1960).(After Pettijohn et al., 1971,fig. 3-14,p. 97;coueesy Springer, New York.) Fig. 3-122.Dependency of the ratio of the permeabilities (vertical (kl)/parallel (kII) to the bedding on kll as illustrated by measurements of the Lias, Dogger, and Valendis sandstones of the Gilhorn Basin in northwest Germany. Near the center, k l approaches the value of 1, while kll increases. (After Riihl and Schmid, in: Von Engelhardt, 1960,fig. 48;courtesy Springer, Berlin.)

chanical and chemical compaction? (2) What are these changes relative to each other? It has been shown by Jobin (1962) that the magnitude of permeability of sandstones controls the localization of uranium precipitation, so that answers to these and similar questions may be of assistance in the investigations of secondary changes of sediments and in exploration for ores. The control of various parameters on porosity and permeability were discussed in numerous other publications, of which a few are considered below. Rogers and Head (1961) offered a composite plot of groups of samples with four different median diameters. Figure 3-123 shows a general decrease in porosity with an increase in sorting coefficient (= an increase in spread of grain sizes). For groups of sands with the same median diameter, the relationship between porosity and sorting coefficient is approximately linear except for very well-sorted sands. The highest porosities in Fig. 3-123

K.H. WOLF AND G.V. CHILINGARIAN

220

approach the theoretical maximum of 46.7% for uniformly sized spheres with cubic packing. Where sorting is very good (i.e., where the coefficient is very low), the curves of different median diameters may coincide. When the sorting becomes poorer, the curves diverge from one another. Also, as the median size decreases, the porosity increases for the same sorting coefficient. In Fig. 3-124,these relationships are more readily shown when the median diameter is plotted against porosity for various sorting coefficients. When the sediment is well sorted, the porosity is independent of the actual grain size and is only a function of sorting coefficient. In cases where the sorting is poorer, the porosity is increasingly dependent on median diameter as well as on sorting. Rogers and Head (1961)also noted that sphericity decreases slightly with decreasing grain size and concluded that the increase in porosity with decreasing grain size is due to a decrease in sphericity and consequent poorer packing. Where all sand grains are perfect spheres, porosity is independent of absolute grain size. Beard and Weyl (1973)investigated the relationship between porosity, permeability, and texture of artificially mixed, dry- and wet-packed sands using eight grain-size subclasses and six sorting groups. Their conclusions can be summarized as follows: (1)In unconsolidated, artificially packed natural sands, the porosity is independent of grain size for sands of the same sorting, but porosity varies with sorting. (2) Wet-packed porosity probably represents the minimum porosity of an unconsolidated, clay-free sand following mechanical rearrangement of the particles before burial. (3)In compaction studies, one should always compare sandstones of the same sorting range, especially when compaction gradients in vertical and horizontal stratigraphic sequences are to be established, because of the variation of porosity 2.2,

1

I

I

I

I

,

1

1

, I

POROSITY, oh

Fig. 3-123. Relationship between porosity and sorting coefficient of sands with various median sizes. A = median diameter = 0.106 mm;B = median diameter = 0.151 mm; C = median diameter = 0.213 mm;fl = median diameter = 0.335 mm. (After Rogers and Head, 1961, fig. 1, p. 469; courtesy J. Sed. Petrol.)

DIAGENESIS OF SANDSTONES AND COMPACTION

221

POROSITY, %

Fig. 3-124. Relationship between porosity and median size of sands with various sorting coefficients. A = S o = -2.086;B=So = - 1 . 6 2 6 ; C = S o = ~ 1 . 2 7 9 ; D = S 0 = - 1 . 1 2 8 ; E = So = "1.061. (After Rogers and Head, 1961, fig. 2, p. 470; courtesy J. Sed. Petrol.)

with sorting. The same applies to many other studies of mechanical and chemical diagenesis. (4) Permeability decreases as grain size decreases and as sorting becomes poorer. (5) Low sphericity and high angularity probably

.. .

I

I

20

I

I

,

30

MEAN GRAIN SIZE, Bunitr

Fig. 3-125. Relationship between the porosity and mean grain size for all sandstone samples of Paluxy Formation, North Texas. (After Dodge et al., 1971, fig. 8 , p. 1826; courtesy Am. Assoc. Pet. Geologists.)

222

K.H. WOLF AND G.V. CHILINGARIAN

increase porosity and permeability. As to a review of the absolute porosity and permeability values in sandstones, as well as for newly obtained values by Beard and Weyl (1973),the reader is referred to their original publication. Dodge et al. (1971)also investigated the various petrographic and directional parameters of Cretaceous, shallow-marine, blanket sandstones that control both porosity and permeability. Mean grain size is plotted versus porosity in Fig. 3-125.These sediments have a high range of both porosity and grain size and lack a cement. There is an obvious absence of any correlation, probably as a result of an insufficient grain-size range in this particular study. Usually, porosity increases with decreasing mean grain size. On the other hand, Fig. 3-126 presents relationship between the permeability and mean grain size of the same sediment and shows a definite correlation: the permeability decreases with a decrease in the mean grain size. In other cases, where cement is present, the data shows a wider spread. The relationship between porosity and percentage of cement is shown in Fig. 3-127.The spread in the area of 0--5.0% cement and 25.0-40.0% porosity may be the result of one or more factors, e.g., variation in the grain size, sorting, and packing. It seems that up to 5% cement may be present before cementation starts to affect the porosity. The cement content is plotted versus the permeability in Fig. 3-128and the spread may be explained by factors mentioned above for Fig. 3-127,e.g., grain size. As shown in both figures, the permeability and porosity decrease with increasing cement content, as would be expected. But again, as mentioned earlier, in detailed investigations of such a

1.0

2.0

3.0

4.0

MEAN GRAIN SIZE, Bunits

Fig. 3-126. Relationship between the permeability and mean grain size for all Paluxy sandstone samples, North Texas. (After Dodge et al., 1971, fig. 9, p. 1826; courtesy Am. Assoc. Pet. Geologists.)

DIAGENESIS OF SANDSTONES AND COMPACTION

01

I

I

I

K)

I

20

223

I

CEMENT CONTENT, %

Fig. 3-127. Relationship between the porosity and cement content for D-1 outcrop, Dexter Sandstone member of the Woodbine Formation of the Gulf Series (Upper Cretaceous), Denton County, North Texas. (After Dodge et al., 1971, fig. 10, p. 1826; courtesy Am. Assoc. Pet. Geologists.)

a m

e

.a 0

2-

t

4 m 4 W

I

a W

a

0

3

10

15

20

25

CEMENT CONTENT, %

Fig. 3-128. Relationship between permeability and cement content for G-6 outcrop, Grayson County, Dexter Sandstone member of the Woodbine Formation of the Gulf Series (Upper Cretateous), North Texas. (After Dodge et al., 1971, fig. 11, p. 1826; courtesy Am. Assoc. Pet. Geologists.)

224

K.H. WOLF AND G.V. CHILINGARIAN

Fig. 3-129. Isometric diagrams showing the variation in cross-bedding dip direction and dip angle and their theoretical influence on permeability. A = low-angle cross-bedding. B = high-angle cross-bedding. Solid arrows indicate maximum directional permeability, whereas dashed arrows indicate minimum directional permeability within a given crossbed set. (After Dodge et al., 1971, fig. 12, p. 1827; courtesy Am. Assoc. Pet. Geologists.)

nature, directional variations in the rock’s mass properties must be considered. As shown in Fig. 3-129,the cross-bedding dip angles in a sandstone may affect fluid flow. If the dip angle is low, the flow will tend to occur along the laminae, whereas if the angle is high, various laminae will have to be crossed by the fluid and the flow will be impeded, depending on the various characteristics of the laminae, e.g., grain size. It should be noted that differential packing, i.e., varying degrees of packing from layer to layer, was not the result of burial in these particular sands, but was the result of the changing transportational and depositional hydrodynamic conditions. Only occasional reference is made in publications to the relationship between porosity and permeability, even though there is a large amount of data available on both porosity and permeability. It seems that quantitative relationships between the two properties should have been given more attention. In the example presented by Dupuy et al. (1963),there is a straight-line relationship when the permeability is plotted on the logarithmic and the porosity on the arithmetic scales, respectively (Fig. 3-130).(See also the data by Shenhav, 1971,given below.) In environmental reconstructions and stratigraphic correlations, it is particularly important to study the grain size, because it reflects the energy conditions. Also, mean grain size may be related to the mineral composition of the sands. Berg and Davies (1968),for example, showed that the quartz content increases with increasing mean grain size of quartz (Fig. 3-131).Also of interest is the relationship between the depositional environment and size and quantity of quartz (Fig. 3-132).In their example, Berg and Davies were able to establish four distinct environmental groups, although for clear field discrimination the knowledge of sedimentary structures and stratigraphic relationships is often required. Inasmuch as both porosity and permeability are a function of the quartz grain size and quartz content, these two parameters should also be related to the original sedimentary environment. This was

DIAGENESIS OF SANDSTONES AND COMPACTION

225

POROSITY, %

Fig. 3-130. Relationship between porosity and permeability for sandstones from Champ de Cazaux (Albian). (After Dupuy et al., 1963, fig. 17; courtesy 6th World Pet. Congr., Frankfurt am Main.)

indeed found to be the case by Berg and Davies (Fig. 3-132).The finest sediments have the lowest permeability, whereas the coarser sandstones have higher permeabilities. Similar results were obtained for porosity. Plots of permeability and porosity versus depth may also reveal vertical changes (through geologic time) in textures, which may occur in sediments of the same type of environment. Inasmuch as the primary, depositional environments control the textures, fabrics, structures, composition, and mass properties of sandstones, which, in turn, may influence subsequent diagenesis, including compaction, then they should be taken into consideration especially in regional and vertical stratigraphic compactional investigations. An extention of the above study can be found in the work of Shenhav (1971).He investigated the petrography, porosity and permeability of Cretaceous sandstone oil reservoirs, and found that the reservoir characteristics are determined by depositional environment and post-depositional changes. Shenhav recognized three main sandstone types each having its own distinct mineralogical, textural and stratigraphic features. Within these sandstones, four main varieties of porosities were defined : (1)Intergranular porosity, with

K.H. WOLF AND G.V. CHILINGARIAN

226

-90-

I-LAGOON

80-

\

X - BEACH OR UPPER SHOREFACE

a- MIDDLE

SHOREFACE I i N c L WASHOVER 5 s )

0- LOWER SHOREFACE

*-

\

I

L

BEACH AND UPPER SHOREFACE

\

(WASHOVER)

70-

\\ *

\

\

A\.

\

\

\

\

\

\

ri090

\

rXy2.096

\

\ I

M

0.05 SILT

0.10 VERY FINE SAND

0.1s

0 20

I

5

FINE SAND

MEAN GRAIN SIZE OF QUARTZ ( i t , r n r n

Fig. 3-131. Relationship between quartz mean grain size and quartz content in Muddy Sandstone, Bell Creek Field, Powder River County, Montana. (After Berg and Davies, 1968, fig 6, p. 1895; courtesy Am. Assoc. Pet. Geologists.)

crystalline cement at aminimum, depends mainly on the packing, sorting and orientation of the grains. There is an absence of pressure solution and the porosity ranges up to 32%. (2) Intercrystalline porosity (i.e., porosity between the crystals of the cement), with relatively high crystalline cement content, reaches a value of up to 8%. Porosity is related to mineralogy of the cement, i.e., the dense, xenotopic calcite cementation gives rise to low values, whereas the hypidiotopic" or idiotopic** dolomite cementation results in larger values of intercrystalline porosity. (3) Intermediate porosity between 1 and 2 ranges from 8 to 20%. The pores are bounded by both crystalline cement and by grains. (4)Intracrystalline porosity results from leaching of the centers in sanidine crystals. Only a small percentage (less than l%), of sandstones have a porosity of this type, which is ineffective in nature. Shenhav, as others before him, found a direct relationship between permeability and porosity; however, in a plot of porosity (arithmetic scale) versus permeability (logarithmic scale), the points have a tendency to spread out over a large area because of variation in the size and shape of grains (e.g.,

* Majority of the constituent crjrstals are subhedral (Friedman, 1965, p. 648).

** Majority of the constituent crystals are euhedral (Friedman, 1965, p. 648).

DIAGENESIS OF SANDSTONES AND COMPACTION

227

50,001

I BEACH AND UPPER SHOREFACE

ro.ooc spoc

500

X

MIDDLE SHOREFACE

u) IOOC

w

-

X

. \

+\

2

2 I

>

t

i

IOC

z

50

4 W

A

LAGOONAL

...

A

10

5

\

AND LAGOO~AL O

I I

O

T

.

SILT

0

5

\

'.

0

1

0.10 VERY FINE

1

' *

0.15

I

0.20

0.25

FINE SeiND

Fig. 3-133). Both porosity and permeability depend on the average grain size, but the effect of grain size on the porosity decreases with decreasing porosity. This is the result of an increase in the cement content. Figure 3-133 presents the general trends (mainly above 8%porosity) of the porosity-versus-permeability curves for various grain sizes. In the case of sandstones with intergranular porosity, the coarser varieties are more permeable than the finer ones of equal porosity. In the case of very fine sandstones with a high clay matrix content, there is no distinct relation between porosity and permeability (Fig. 3-135).Shenhav plotted porosity versus permeability for each one of the sandstone types, as exemplified by Fig. 3-134for type I. He

K.H. WOLF AND G.V. CHILINGARIAN

228

HORIZONTAL PERMEABlLiTY, M D fLOG

SCALE)

Fig. 3-133.Relationship between porosity and horizontal permeability for the various sandstone types of Helez Formation, Israel. (After Shenhav, 1971, fig. 25, p. 2222; courtesy Am. Assoc. Pet. Geologists.)

001

01

OD

I

0

10

w

mo

100

HORIZONTAL PERMEABILITY, M D (LOG

SCPLE)

Fig. 3-134.Relationship between porosity and horizontal permeability of Type-I sandstone samples of Helez Formation, Israel. (After Shenhav, 1971, fig. 22, p. 2219;courtesy Am. Assoc. Pet. Geologists.)

229

DIAGENESIS OF SANDSTONES AND COMPACTION

HORIZONTAL PERMEABILITY. M 0

(LOG

SCALE1

Fig. 3-135.Plot of porosity versus horizontal permeability trends of Helez Formation sandstones having various grain sizes. (After Shenhav, 1971, fig. 21, p. 2219;courtesy Am. Assoc. Pet. Geologists.)

also presented a composite curve showing the fields of data distribution for various sandstone types (Fig. 3-133). According to Shenhav, petrographic considerations allowed the recognition of three different depositional environments, giving rise to offshore marine, tidal channel or lagoonal, and dune sandstones. These sandstones had the characteristic mass properties shown in Table 3-XXXIII. The amount TABLE 3-XXXIII Mass properties of various types of sandstones (after Shenhav, 1971, figs. 22-25, pp. 22192222) Type of sandstone

Dune (eolian coastal)

Mass properties porosity (%)

permeability (md)

24 (avg.) up to 32

200 (avg.), max. over 2000 50 (avg.) up to 2000 less than 30

Tidal channel and/or lagoonal

16 (avg.) up to 30

Offshore marine

very low up to 16

Reservoir characteristics good intermediate poor

230

K.H. WOLF AND G.V. CHILINGARIAN

and mineralogy of cement, which are determined by the sedimentary and diagenetic milieu, are the main factors controlling the porosity, whereas permeability is related only to the amount of cement. In the case of intergranular porosity, permeability depends on the average grain size. One of the well-known difficulties confronting sedimentologists and petrologists is the large amount of time involved in the determination of precise data of the various petrographic variables by employing instrumental techniques. In many instances, despite a reduction in precision and reproducibility, as well as possible variations between researchers (or “operators”), useful methods have been proposed to speed up the procedure for petrographic work in both field and laboratory investigations. The charts for quick determinations of roundness and sphericity of sand grains and pebbles have been available for some time, in addition t o the charts for visual percentage estimations. Beard and Weyl (1973)have presented a new set of visual textural comparators, i.e., eight sets of photographs of sands ranging from very fine to very coarse, each composed of six photographs depicting extremely good to very poor sorting range. In combination with the other charts mentioned above, they permit rapid estimation of a total of four textural variables. Beard and Weyl suggested that these charts are particularly useful in describing compacted and silica-cemented sandstones. Upon considering some examples of investigations on the porosity and permeability variations that appear to have been controlled by primary depositional factors, one should examine the relationship between the depth of burial and mass properties of sedimentary rocks. Maxwell (1964)has stated that it is a valid generalization that porosity varies inversely with depth of burial and, less certainly, with geologic age. These conclusions are well supported for shales. Although the absolute values differ, porosity of shales decreases rapidly in the first few thousand feet of burial and much more slowly thereafter (see Rieke and Chilingarian, 1974). Less information is available on the variation of sandstone porosity with depth, because of the extreme variability in various parameters. Even less information is available on the comparative relations between porosity and depth of burial for interbedded sandstones and shales of the same basin or stratigraphic section. Figure 3-136 shows such a comparative approach by F’roshlyakov (1960). The porosityaepth relationship for shale is better defined than that for the sandstone. Maxwell stated that neither the shape nor the trend of the sandstone curve could be anticipated from the published information. He collected data on the porosities of natural quartz sandstones and compared them with published data. Inasmuch as temperature is an important variable affecting the compaction of quartz sands, he also collected temperature data and compared it with porosities in natural sandstones. Although there are certain difficulties in obtaining accurate temperature data, and in most wells

DIAGENESIS OF SANDSTONES AND COMPACTION

2000

231

c

* = SAND 0

= SHALE

POROSITY '10

Fig. 3-136. Relationship between porosity and depth o f burial for Jurassic-Lower Cretaceous elastic sediments (sands and shales), Cis-Caucasus, U.S.S.R. (After Proshlyakov, 1960; in: Maxwell, 1964, fig. 1, p. 698; courtesy Am. Assoc. Pet. Geologists.)

the temperature gradient has not been measured, enough information is available, however, to show some significant relationships among porosity, depth, and temperature. Maxwell discussed eight examples, of which one is reproduced here (Fig. 3-137). The porosity data has been plotted as bar graphs, each bar summarizing the measurements of one individual sandstone section several tens of feet thick. The small solid rectangles indicate the average porosities. Each figure summarizes the data for sandstones of the same geologic province and approximately the same age t o enable comparison between the depth of burial and temperature variations. Figure 3-138 is an example where the decrease in both maximum and average porosities with depth is striking. Also, the change of maximum porosity is nearly linear with depth, which is distinctive for these particular sandstones. It should be pointed out, however, that Maxwell discussed some exceptions to that rule. Maxwell (1964) also reported on the results of experiments on the compaction and cementation of quartz sands under conditions simulating deep burial. Results of the short-time compression tests at room temperature suggest that high porosity can persist to great depths. At pressures simulating a depth of 40,000 ft, a 25% porosity may remain, and if the normal hydrostatic pressure is present, the porosity may be as high as 30%. As Maxwell pointed out, experimental conditions do not necessarily resemble natural

232 I5

5 1

I

I

I

25 I

I

i1 3

5

K.H. WOLF AND G.V. CHILINGARIAN ~

.

~

5.000

1

10.000

20.000

/

-

15.000

/

/

/

/

ym, /

/

l5.000

A

20.000

POROSITY,%

Fig. 3-137.Relationship between the porosity: and depth of burial for Pennsylvanian and for two shallow Permian sandstones (at about 3000 and 5000 ft), North Texas and Oklahoma. Porosities from depth of below 20,000 ft calculated from sonic and neutron logs. Data from Shell Oil Company, Socony-Mobil, and from f i l l and Taliaferro (1949). Small rectangles give average pay porosity, whereas figures in parentheses represent number of porosity determinations. Temperature gradient in "C/lOOOf t is noted along dashed line limiting maximum porosity. (After Maxwell, 1964, fig. 5, p. 701; courtesy Am. Assoc. Pet. Geologists.)

conditions of compaction, because many variables are not considered in the laboratory experiments. In his tests, Maxwell considered composition of the particles and pore fluids, temperature, overburden pressure, pore-fluid pressure, and time; however, the effects of grain size and sorting, for example, were not evaluated. Maxwell presented a summary of porosity-versus-depth data for the relatively pure, well-sorted quartz sandstones as shown in Fig. 3-138,which is based on figs. 4, 5, 7, 9 and 10 of his 1964 publication. Two experimental curves based on the experimental data published by him in 1960 are superimposed in this figure. Noteworthy is the variability of the curves for the Oligocene sandstones as the thermal gradient changes from 7"--1O0C/10O0 ft. The following observations can be made on examining Fig. 3-138:(a) the Oligocene sandstones have larger maximum porosities at

DIAGENESIS OF SANDSTONES AND COMPACTION DEPTH

oo oO

I

MAXIMUM POROSITY X 5

10 10

I5

20

25

30

35

40

45

1

5-,000

5.000

10.0001 10.000

15.000

- 15.000

20.000

- 20.000

25.000

26.500'

- 3o.OOo

30.000

35.0%

233

5

20

25

30 ..

. .

4.0 -

a5 ._

35'W

Fig. 3-138. Relationship between the maximum porosity and depth of burial for various sandstones. (After Maxwell, 1964, fig. 12, p. 706, based on figs. 4, 5, 7, 9, and 10 of Maxwell, 1964, and experimental data of Maxwell, 1960, figs. 6 and 11; courtesy Am. Assoc. Pet. Geologists.)

shallow depth than the Miocene ones, but the two depth-versus-porosity curves cross at about 13,000 ft, i.e., the porosity of Oligocene sandstones decreases more rapidly with depth than that of Miocene sandstones; (b) the Ordovician sandstones have larger maximum porosities than the much younger Pennsylvanian sandstones. This behavior can be predicted by using the equation developed by Maxwell (1964, p. 702) from the experimental data for quartz sand (see also description of Fig. 3-139). In Fig. 3-139, two sets of curves were plotted, i.e., curves of the observed data taken from Fig. 3-138 and curves based on the calculated data. Although the two sets of curves agree only in a general way, the calculated data enable prediction of the actual field observations to a certain degree. The calculated depth at which the Oligocene and Miocene sandstones are expected to have identical porosity values depends on the values of the assumed original depositional porosity for each sandstone and, therefore, on the slopes of the two curves. The porosity values in Fig. 3-139a are anomalous, because extrapolation of the curves to zero depth gives improbably large initial maximum porosity values of over 50%for both sandstones. Maxwell (1964) offered two possible explanations: (a) measurements are too high as the sediments were poorly consolidated and the samples were disturbed or (b) one could assume that initial porosities of over 40% were maintained to depths of 7,000-8,000 f t because of lack of compaction. In view of the straight-line character of the

K.H. WOLF AND G.V. CHILINGARIAN

234

5,000

-I

t

0

10

20

30 POROSITY,%

10

20 _.

5,000

30 .. J C M. ‘63

Fig. 3-139.Relationship between the maximum porosity and depth of burial for various sandstones. a. Observed curves of maximum porosity obtained from Fig. 3-138;b. porosi~-~ where Gi = ties calculated by means of the equation ( @ i / q ! ~ f )=~ t[e(T-540)/(0-0186T] porosity at time of deposition, @f = final porosity, t = time in days, T = temperature, degrees Kelvin. This equation is based on experimental data and predicts porosity of sandstones for simulated “depth” of 26,500 f t (Maxwell and Verrall, 1954).Slope determined by extending line from this point toward assumed value of initial porosity, @i. (After Maxwell, 1964,fig. 13,p. 707; courtesy Am. Assoc. Pet. Geologists.)

maximum porosity boundary curves, the latter explanation seems unlikely. As Maxwell pointed out, however, the available data are inconclusive. In the experiments reported by Maxwell (1960, 1964), compaction of the pure quartz sand occurred mainly by fracturing of grains, accompanied by rotation of grains and fragments. Silica cementation was noticeable, but its amount was very small. Both physical yielding and chemical precipitation of cement are involved in the compaction and induration of natural quartz sandstones. In cases of the presence of largest porosity, it can be assumed that the sandstone unit has undergone least amount of secondary quartz cementation. Under these conditions, compaction may occur by physical failure of grains, similar to that observed in the experiments, in addition to the occurrence of pressure solution. One should note here, however, that many researchers have suggested that such fracturing occurs only in the shorttime laboratory experiments in which the long geologic time, available during the natural compaction in most sedimentary basins, is not considered. In the presence of abnormally high pore pressures, the graiwto-grain pressure may be reduced sufficiently to diminish compaction so that abnormally large

DIAGENESIS OF SANDSTONES AND COMPACTION

235

porosities could persist to great depths. Although the available data is inadequate for reaching any definite conclusions, Maxwell suggested that abnormally high pore pressures do not guarantee persistence of unusually large porosities a t depth, and stated that this problem offers a fertile area for more intense research. Maxwell (1964)discussed the cause of linearity of his porosity-versusdepth curves. This linearity was quite unexpected, as theoretically one would expect a change in rate of reduction of porosity with depth. Maxwell, therefore, offered an explanation for the linearity on the basis of experimental results obtained by various investigators: (1)At room temperature and a gage pressure of 1atm, the short-time compaction experiment showed porosity reduction per unit of load which was largest at the small initial loads and diminished rapidly with progressive reduction in porosity under increasing pressure. The plotted results gave a porosity-“depth” curve with a shape similar to that of the shale curve presented in Fig. 3-136. (2) When both the temperature and “depth” (= pressure) were held constant, porosity first diminished rapidly during the first few days and then decreased at a decreasing rate throughout the experiments which lasted up to 100 days. (3)When dry quartz sand was tested at a pressure of 1atm, the grains became progressively and distinctly weaker as the temperature was raised. The physical failure, therefore, was shown t o be strongly dependent on temperature. (4)When sea water and oil-field brines were surrounding the grains, failure of the grains was again strongly temperature dependent; however, porosities were lower than in the experiments performed under conditions described under (3).The solutions seemingly caused an additional weakening of the quartz grains. Pressure solution was negligible. (5)The solubility of quartz in distilled water increases with increasing temperature, especially at higher pressures (Kennedy, 1950).On the other hand, the dissolved chemical components in sea water and brines have little effect on the solubility of silica (Krauskopf, 1950;Siever, 1962). The solubility of quartz in natural waters, therefore, should increase as temperature increases. From the above-presented information, it appears that: (a) compaction and consequent porosity decrease in quartz sands should proceed at a decreasing rate with burial if the temperature is presumed to be constant; (b) on the other hand, the increase in temperature will cause a loss of porosity with burial at an increased rate, due t o both decreasing physical strength of the grains and as A result of increasing solubility of the quartz grains. Maxwell then suggested that these two trends tend to balance one another, so

K.H. WOLF AND G.V. CHILINGARIAN

236

that the curves of maximum and average porosity may approximate straight lines. The tests performed by Maxwell indicated that compaction continues as long as porosity is present; there is no indication of an equilibrium porosity for a particular depth which would persist throughout geologic time. Teodorovich and Chernov (1968) developed a relationship between the depth of burial, D, in m and permeability, h, in md for sandstones at a depth range of 500-4500 m: log k = 2.803 X e4*000074Dor (approx.) log k = 2.88 - 0.0002D

(3-4)

For siltstones at the same depth range the formulae are: log h = 2.961 X e4.000147 or log k = 2.87 - 0.003D.

(3-5)

Variation in the permeability of sandstones is much greater than that for siltstones at the same depth (see Table 3-XXXIV), because of the higher heterogeneity of the former, e.g., at a depth of 3950-4450 m, permeability varies from 30 to 350 md and, sometimes, up to 800 md (average k = 114 md). When the cement content is less than 7%, porosity at a depth of 4250 m is higher than 20-25% and permeability is higher than 150-200 md. With increasing depth of burial from 4250 to 6000m, permeability decreases 2.2-2.4 times for siltstones and 1.5-1.8 times in the case of sandstones. Meade (1966) observed that variations in the porosity of sands and in their water content are not predictably related to depth and an increase in the overburden load. Figure 3-140 shows this for a selected group of sediments that were exposed to their approximate maximum overburden loads at the time of sampling. All these sediments are Cenozoic in age and most are inorganic and terrigenous clastics, except for sample I11 containing about 10% siliceous skeletal material, sample I1 which is clayey, and sample IV with 10-20% CaC03. The curves are drawn in such a way that the area TABLE 3-XXXIV Variation in porosity and permeability with depth for sandstones and siltstones of A p sheron oil- and gas-bearing province in U.S.S.R. (after Teodorovich and Chernov, 1968, P. 88) ~~

Depth (m)

650 4250 5000 6000

Sandstones

Siltstones

porosity (%)

permeability (md)

porosity (X)

permeability (md)

26 f 4 17 i 4 15.5f 4 14 f 4

460 f 110 i 87 i 63i

28.5 * 16 i 14.7 * 13 i

5 0 0 i 75 4 0 i 10 26+ 8 17 f 5

80 25 20 20

4 4 4 4

DIAGENESIS O F SANDSTONES AND COMPACTION

237

DEPTH OF BURfAL IMI

Fig. 3-140.Relation between porosity and depth of burial in meters in selected clays and claystones (A) and selected sands and sandstones (B). I = Recent, Lake Mead on Colorado River (Gould, 1960, p. 176, lower 3 graphs); II = Recent, Santa Barbara Basin off southern California (Emery and Rittenberg, 1952, p. 755.);III = Recent and older (?), western Bering Sea (Lisitsyn, 1956, 1959,fig. 16);IV = Recent and older (?), eastern Black Sea (data from Ostroumov and Volkov, 1964,pp. 94-95); V = Recent and older (?), continental slope off Nova Scotia (Richards and Keller, 1962); VZ = Recent, Orinoco River delta (Kidwell and Hunt, 1958,p. 808); VII = Pliocene and Pleistocene, central California (Meade, 1963a;curves adjusted for artesian pressure); VIZI = Pliocene to Recent, Baku Archipelago (Koperina and Dvoretskaya, 1965, fig. 1; plus data from Korobanova, Kovaleva, Kopylova, and Safokhina, 1965,pp. 128-130); I X = Tertiary, Venezuela (Hedberg, 1936,p. 256);X = Miocene and Pliocene, Po Valley (Storer, 1959, p. 523); XI = Miocene, southern Louisiana and southeast Texas (Maxwell, 1964, p. 704).Where bulk density or water content (by weight) were reported, porosity was computed assuming a particle density of 2.60g per cm3. (After Meade, 1966,fig. 1,p. 1086;courtesy J. Sed. Petrol.)

above the lines gives the volume of solid particles, whereas the area below represents the interstitial volume. As to the influence of particle size on water content and porosity (the two are numerically equal because both are expressed as percentage of bulk volume), Meade stated that particle size may be as important a control as overburden load on porosity, especially during the early stages of compaction. A comparison of Figs. 3-140Aand 3-141A shows that an order-of-magnitude difference in diameter between 0.001 and 1 mm may cause about the same amount of porosity variation as an order-of-magnitude difference in depth of burial. With an increase in pressure, the finer sediments are compacted more rapidly than the coarser ones, and the relationship between grain size and porosity becomes less pronounced (Fig. 3-141B). A significant influence of grain size on porosity is still apparent at overburden loads approaching 100 kg/cm2. For the sediments presented in Fig. 3-141B, the magnitude of change in porosity with changing grain size is about equal to that related to depth of burial (cf. the curves in Fig. 3-141B with curve VII in Figs. 3-140A

K.H. WOLF AND G.V. CHILINGARIAN

238

50-

5

1

P MEDIAN DIAMETER fMMJ

Fig. 3-141.Relations between porosity and median particle diameter in selected Recent surface or near-surface sediments under overburden loads less than 1 kg/cm2 (A), and in sediments under overburden loads between 7 and 70 kg/cm2 (B), I = Lake Mead on Colorado River (Sherman, 1953, p. 399);ZZ = Lake Maracaibo, Venezuela (Sarmiento and Kirby, 1962, p. 719); IZZ = reservoirs in western United States (Hembree, Colby, Swenson, and Davis, 1952,p. 39);ZV = Gulf of Paria (van Andel and Postma, 1954, p. 108); V = North Sea (Fuchtbauer and Reineck, cited by Von Engelhardt, 1960,p. 15); VZ = continental shelf off southern California (Hamilton and Menard, 1956,p. 757); VZZ = San Diego Bay and adjacent continental shelf (data from Shumway, 1960,pp. 454-457); VZZZ = Rivers in Japan (Komura, 1963,p. 266);ZX = Pliocene and Pleistocene alluvium, central California (Meade). (After Meade, 1966, fig. 2, p. 1087; courtesy Am. Assoc. Pet. Geologists.)

and 3-140B). Another group of sediments in California exposed to overburden pressure of 5-60 kg/cm2 (Meade, 1963b), demonstrated that their downward decrease in particle size (together with changes in other factors) reversed the expected effect of increasing load and caused a systematic downward increase in porosity. With a further increase in pressure as a result of overburden, an inverse relation between the grain size and porosity changed to a more direct correlation. According to Meade (p. 1087), “the more rapid decrease in the water content of clays eventually overtakes the slower compaction of sands, perhaps at some depth of burial near 1km”. This is supported by the observations that deeply buried coarser sediments and their equivalent rocks are commonly more porous than adjoining claystones and shales. In his section on compaction of sands, Meade (p. 1096) stated that most sands are only slowly compacted during the early stages of compaction; however, relatively few unequivocal data is available to demonstrate this, because of the difficulty in collecting unconsolidated sands without disturbing them. Curve VII in Fig. 3-140B represents silty sands, and it seems that the silt gave the sand enough cohesion to withstand coring and sampling without too much distortion. Curve X I in Fig. 3-140B represents quartzose sandstones in approximate equilibrium with their present overburden pres-

239

DIAGENESIS O F SANDSTONES AND COMPACTION

sure. During the early stages of compaction, the parameters that control porosity are mainly textural, i.e., grain size, sorting, roundness, shape, and flexibility of certain types of grains (Von Engelhardt, 1960,pp. 3-16; Fraser, 1935; Gaither, 1953; Hamilton and Menard, 1956). The results of experiments on the influence of sorting and roundness on compaction of pure quartz sands are given in Figs. 3-142A,B.In the pressure range up to 100 kg/cm2, compaction occurs as a result of rearrangement of sand grains into a more dense packing system and, to a minor degree, by elastic compression of the individual grains. At pressures above 100 kg/cm2, compaction increased as a result of cracking and shattering of the grains in the compression apparatus. In agreement with other investigators, however, Meade doubted whether this cracking and shattering would occur under natural conditions of com-

i

A -

50i u.

- 0 1

t

0.4

10

100

1000

1

10

3 1000

100

: 4oor C

50

1-

MICA

20%

I

PRESSURE

1

to

IKG PER CM'J

Fig. 3-142.Influence of different factors on relations between porosity and pressure in sands, as determined in laboratory experiments. A. Influence of sorting in well-rounded quartz sands (Roberts and de Souza, 1958);sorting index (a@)defined by Inman (1952, pp. 135-136); median diameter of two better sorted sands = 0.60 mm; median diameter of sand with poorer sorting = 0.48 mm. B. Influence of rounding of quartz sands, 0.42-0.84 mm in size (Roberts and de Souza, 1958). C. Influence of mica particles mixed in different proportions with rounded quartz sands (Gilboy, 1928,p. 560);particPes of both constituents are 0.42-0.59 mm in size. (After Meade, 1966,fig. 10,p. 1097; courtesy A.m. h s o c . Pet. Geologists.)

240

K.H. WOLF AND G.V. CHILINGARIAN

paction of sands. The slow rate of pressure increase in nature, extending over thousands and millions of years, permits other processes to be operative, such as plastic flow, pressure solution, and reprecipitation. Sands that are composed of minerals other than quartz may respond differently to pressure, with the softer grains being more readily deformed. That the greater porosity of well-sorted sands, in contrast to more poorly-sorted ones, persists during early compaction is indicated by the results presented in Fig. 3-142A. The influence of particle roundness is demonstrated in Fig. 3-142B, where angular sands show greater initial porosities. This reflects the instability of the initial packing of the angular grains. The angular particles are more compactible than rounded ones of the same grain size. The pronounced influence of platy and flexible mica particles on the behavior of sands in compaction experiments, is demonstrated in the five graphs of Fig. 3-142C. The porosity, compressibility, and elasticity of sand increase with increasing mica content. Experiments by McCarthy and Leonard (1963) support these results and suggest further that the finer the mica flakes among the sand grains, the greater the increase in porosity per unit increase in mica content. It should be noted, however, that a permanent, rather than elastic, deformation of mica plates may also take place during the early stages of compaction of micaceous sands. At pressures between 0 and 100 kg/cm2, grain size is the most important parameter influencing compaction. Not only is it inversely related to water content and porosity of the sands, but the size of the particles influences most other factors that control mechanical and chemical compaction. The effect of grain size may be so strong that the expected decrease in porosity with depth of burial may be obscured. The presence of mica may have a greater influence on the mass properties of sands than textural variations, such as those of rounding and sorting, of the non-mica grains. It is important to stress here, that more complex combinations of sand, silt, and clay should be experimentally examined in order to determine the influence of their relative proportions on the rate and degree of compaction. In many laboratory investigations on the changes of mass properties with increasing pressure (simulating burial of sediments in basins), the possible effects of temperature were not taken into account. It is, therefore, of particular interest to note that Somerton and El-Shaarani (1974) found that the compressibility of sandstones increases with increasing temperature. The effect of temperature was more pronounced at the lower effective pressures. Sawabini et al. (1974) determined relationship between the void ratio (= volume of pores/volume solids) and effective pressure, pe (Pe = pt -pp, where pt is the total overburden pressure and pp is the pore pressure) for unconsolidated sandstones from .a depth of 3000 f t (Fig. 3-143). In the effective pressure range of (F3000psi, the void ratios varied from 0.85 to

DIAGENESIS OF SANDSTONES AND COMPACTION

I

O m

L

,

1

1

1

1

1

HXK)

I

I

1

L

I

24 1

I

QW

EFFECTIVE PRESSURE, psig

Fig. 3-143. Experimental relationship between void ratio and effective pressure for unconsolidated sandstones. (After Sawabini et al., 1974, fig. 9, p. 136; courtesy SOC.Pet. Eng. AIME.)

0.19. These authors tested unconsolidated, medium- to fine-grained, arkosic sand cores obtained from oil-producing formations of Pliocene and Upper Miocene age in the Los Angeles Basin, California. The overburden (external) pressure was held constant at 3000 psi in a hydrostatic (three principal stresses are equal) compaction apparatus, at a temperature of 140"F,while producing the interstitial fluids and thus reducing the pore pressure and increasing the effective (grain-to-grain) pressure. It is the latter stress that causes compaction and, consequently, the one that should be used in plotting porosity-versus-pressurecurves whenever possible. PLASTICITY, COMPRESSIBILITY, DENSITY, AND THIXOTROPY OF SANDY SEDIMENTS AND SANDSTONES

A number of physical and chemical properties of sandy sediments are related to compaction in that (a) the degree and rate of compaction controls these properties, and (b) vice versa, these properties themselves influence the style, rate and degree of compaction. The most important properties referred to here are plasticity, compressibility, density and thixotropy* of sandy sediments (with silt and/or clay), sands, and sandstones. These topics are only briefly discussed in this chapter and only some selectively chosen mate-

* Thixotropy is the ability to gel (become firm) upon quiescence and to become fluid upon agitation.

242

K.H. WOLF AND G.V. CHILINGARIAN

rial is presented. From the outset it should be pointed out that to confine the discussion to sand-sized components is really meaningless, because the four properties listed above are a direct function of purity, or impurity, of the sediments, i.e., the amount of clay and silt admixed with the sandy constituents is extremely important in determining the behavior of the sediment. To extrapolate this argument, deflocculated clay minerals may fall into the group of clay-sized components, although the largest particles are up to 5 p in size and may fall into the silt-sized group. Once the clays are flocculated and form larger aggregate particles, however, they may belong to the silt- and sand-sized classes. It should be remembered, therefore, that even though this book is devoted t o sands, the data on clays given below aids in better understanding the behavior of clayey siltstones and clayey and silty sandstones. This has been pointed out also in Chapter 2. The bulk density* of sedimentary rocks depends on: (a) the density of individual types of minerals present; (b) the proportions of different minerals with varying densities; (c) the porosity; and (d) the amount of fluids in the intergranular pores. Inasmuch as the porosity depends on the degree of compaction and cementation, for example, one can speak of original and secondary densities of the sedimentary rocks. The densities usually increase with increasing degree of compaction and lithification, and details are given below. Figure 3-144presents the hydraulic equivalents of light and heavy minerals commonly found in sediments. In a general way, particles with similar hydraulic properties have a tendency to accumulate together t o form sedimentary laminae or beds as a result of both shape and size sorting, during transportation to and within the depositional environment. The actual density of sedimentary particles is, of course, important, but inasmuch as the shape of the grains controls their behavior during erosion, transportation and settling in fluid media (i.e., water and wind), it is the combination of both density and shape plus size that determines the hydraulic properties of particles. To be able to compare these properties, hydraulic equivalents, like those in Fig. 3-144,have been determined. Although a number of variables control the density of rocks, particular

* Bulk density of sediments has been defined as the weight of the sediment per unit of bulk volume (bulk volume, V , = pore volume, V,, + volume of solid grains, V,) or the mass of the sediment per unit of bulk volume. The mass p is attracted by the earth with a force (weight) having magnitude p g , where g is the gravitational acceleration. For example, pure water, which has a specific weight of 62.4 lb/cu ft, has a density of 1.94 (= 62.4/32.2) slugs/cu ft (gravitational acceleration = 32.174 ft/sec/sec). If density is reported in g/cm3, then it is equal to specific gravity.

243

DIAGENESIS OF SANDSTONES AND COMPACTION

Hobit Aciculor

I [Gypsum IUI

Cordierite

Prisrnotic

[Sillimonite ApOaTourmaline

e-

zircon

Rutilei

Equant

To bu lo r

I

Anotose Feldsporbmmui

Barite Kyonite

Andolusite UIlU mnmnmmmChlorlte mTopoz BIOtiternmmmm mnmmOChloritoid mum Mvscovite ontrnorillonite Im Dlospore olinite

Platy

Density

20

25

30

35

40

Brookite

45

Hematite Ill

5.0

5

Fig. 3-144. Diagrammatic representation of hydraulic equivalence of the commoner light and heavy minerals. (After Griffiths, 1967b, fig. 10-5; copyright 0 1967 McGraw-Hill, New York.)

lithologic units have been found to have a mean bulk density that can vary distinctly from unit to unit (cf. Fig. 3-145, for example). Before entering the subject matter of plasticity, compressibility and re-

Oriskany

Venango

Bradford Miocene

Sespe

Fig. 3-145. Mean bulk'density for rock types, (After Griffiths, 1967b, fig. 18-3; copyright

@ 1967 McGraw-Hill, New York.)

K.H. WOLF AND G.V. CHILINGARIAN

244 Inviscid fluid

- - -

Viscous 1Iuid

IPascalian)--IStokesiani

/I

Visco-elastic fluid IMaxwelii

Elastlco-vIscOus solid (Kelvin)

Elastic solid Rigid solid IHookeanl-IEuclidea~)

Plastics IBoltzrnannl Plastics ilinghaml

Fig. 3-146. Rheological classification of materials. (After Fredrickson, 1964, fig. 55; courtesy F’rentice-Hall, Englewood Cliffs, N.J.)

lated phenomena, it should be pointed out that these concepts imply flow under pressure. The field of rheology is concerned with such phenomena and includes a whole range of “fluid” and “solid” types of materials, as shown in the classification in Fig. 3-146.The flow phenomena as related to one particular type of geologic or natural system, i.e., muddy sediments including sandy deposits, are presented in Fig. 3-147.A comparison between the flow behaviors of Newtonian (viscous) (called Stoke&an fluids in Fig. 3-146)and Bingham fluids is presented in Fig. 3-148. D i spers ions

Shearhardening di spers ionS

V i s c o s i t y dependent on r a t e o f shear

V i s c o s i t y independent of r a t e of shear

O i ~ p e r s i o n sw i t h v i s c o s i t y dependent on tinw o f r e s t and time of shear

D i spe r s Ions w i t h Y Is cos It y independent of time o f r e s t and t i m e of shear

False-body dispersions

Rheopectic d i spe rs i oils

Thixotropic dispersions

Plastic d I spe r s ions

Pseudo-plastic d i s p e r s ions

Diiatant dispersions

Non- rheopect ic

dispersions

Fig. 3-1 47. Classification table summarizing the relationships of the flow phenomena in muddy sediments. (After PryceJones, in: Boswell, 1963; courtesy W.H.Heffer and Sons, London.)

DIAGENESIS OF SANDSTONES AND COMPACTION

I

A

245

B

Fig. 3-148. Schematic diagram of flow behavior of Newtonian (viscous) ( A ) and Bingham ( B ) fluids. Below point 1 on curve A , the flow is viscous, whereas above it the flow is turbulent. Pressure p t is necessary to start the fluid movement (yield point). 1 ’-2 ’: region of plastic flow; above point 2 ’, the flow is turbulent.

Newtonian fluids are the simplest fluids from the standpoint of viscosity considerations, because the viscosity coefficient, p , remains constant at all rates of shear. Viscosity 1.1 is equal to K ( d q / d p ) , where K is a constant and d p / d q is the slope of the straight-line portion of curve A in Fig. 3-148. Curve B in Fig. 3-148illustrates the flow behavior of Bingham fluids. As shown in this figure, a finite, minimum pressure pt must be applied to start the movement of the fluid and is referred to as the yield point. The straightline portion of curve B represents the plastic flow region (line 1’-2’). The ~ equal to K ‘ ( d q / d p ) ,where K‘ is a constant and dqldp is plastic viscocity 1 . 1is the slope of the 1’--2‘ line. Some of the aspects related to the flow phenomena of sediments is determined by their rheology or rheologic properties. These properties, which determine the flow of materials in general, have been investigated in detail by physical scientists and engineers, but only occasionally have been applied to geologic problems (e.g., Boswell, 1961;Elliston, 1963;McNeill, 1963, 1966). The type of flow phenomenon in which the change from a rigid to a fluid condition is produced under ordinary temperature by mechanical action, without the application of heat, is called “thixotropy” (part of the encompassing field of rheology). After being undisturbed for a while, the material will set again to a jelly-like mass, in some cases at once and in others slowly. A water-plus-clay system can change from a gel to either a sol or a more fluid gel as a-response to mechanical disturbance, and then change back to gel when left at rest.

K.H. WOLF AND G.V. CHILINGARIAN

246

As Boswell (1963),for example, has pointed out, differences in texture (angular vs. rounded grains, grain-size distribution, etc.) and composition (e.g., percentage and types of clays in the matrix of sands) result in variations in the rheologic, especially thixotropic, properties of sediments. Differential compaction may depend on these properties, because the capacity to retain moisture and the tensile and shear strengths determine the sediments’ ability to change in response to overburden load. Many intraformational structures are evidence of this, although the present knowledge of rheological properties of sediments in controlling diagenetic disturbances and reworking of unconsolidated deposits is only in its infancy (e.g., Elliott, 1965). Figure 3-149A1is the more commonly accepted way of representing the composition of sediments, i.e., the interstitial water is ignored. If one wishes to consider the rheological properties of sediments, however, then the presFigure ence of water has to be considered, as shown in Fig. 3-149A2,B. 3-149A2also illustrates the changes in properties from “permanent fluid”, through “thixotropic”, to crumbly, plastic and cloddy, depending on the composition. Figure 3-150shows the rheotropic zones in relation to depth below the surface, as one might expect in a sedimentary basin, as a function of grain size. This diagram by Boswell (1963)has been applied to practical geologic problems by McNeill (1963,1966). Below the surface of sediment accumulation (i.e., “belt of variables” or zone of physical, biological, and chemical variables), the following zones succeed each other: (1)permanently fluid muds (clay and/or silt); (2) thixotropic sediments; (3)plastic deposits; (4)“indurated” and compacted sediments. In the thixotropic zone, enough water is present in the sediments to allow them t o flow under shock or repeated vibration. The sediments, however, become firm again when the disturbance ceases. The plastic sediments can flow under stress, but as soon SAND

SILT

WATER

SILT

and

SAND

CLAY

Fig. 3-149A. Diagrammatic representation of varying composition and characters of noncalcareous muds. A1 = dry; A2 = wet. (After Boswell, 1963, fig. 1 ; courtesy W.H. Heffer and Sons, London.)

247

DIAGENESIS OF SANDSTONES AND COMPACTION

SAND and SILT

Fig. 3-149B. Diagram showing relationships of calcareous and argillaceous muds. (After Boswell, 1963, fig. 2; courtesy W.H.Heffer and Sons, London.)

as the pressure is released, the material reverts to a rigid state. The formation

of small faults may be the result of plastic flow. The tendency for creation of fractures, faults, and brecciation increases with depth towards the indurated zones and towards areas where sediments are well cemented. Cementation has been ignored in Fig. 3-150 to simplify the presentation. It is edaphic deposits Belt of Variables

CLAY

I

SILT

hypedaphic deposits

I

0mm

SAND

Fig. 3-150. Diagrammatic representation of diagenetic zones in relation to depth below sea floor and grading. The Belt of Variables and edaphic and hypedaphic deposits extend across the figure. (Cf. also McNeill, 1963, in Elliston and Carey, 1963; and McNeill, 1966.) (After Boswell, 1961, p. 100, in Boswell, 1963, fig. 5; courtesy W.H. Heffer and Sons, London.)

248

K.H. WOLF AND G.V. CHILINGARIAN

significant to consider Boswell’s (1963)statement that compaction of a stratigraphic section of sediment will not proceed evenly with geologic time, but in stages when rheotropic yield values are reached for each sedimentary deposit as burial continues. The resistance to deformation is related to the various index properties of the sediments (including rheotropy or thixotropy, for example), and inasmuch as these properties change from one sedimentary unit to another, compaction will proceed irregularly. In some instances, older beds may require longer time than the overlying younger ones to reach maximum compaction. Sposmatic differential compaction or settling may, therefore, be a normal feature of mechanical diagenesis. As Boswell pointed out, there are various ways in which sediments can offer resistance to stresses and, consequently, the style of the disturbance can vary accordingly. Four examples of this are presented here: (1)“Shearrate blockage” (i.e., dilatancy) in the case of clean, well-sorted, and wet sands, silts, and muds, when the disturbance occurs rapidly. This leads to the formation of a more dense and rigid mass of sediment accompanied by the expulsion of water. As Boswell pointed out (1963,p. 107), originally bedded or laminated sedimentary units that underwent “shock” during an earthquake, when the deposits were still thixotropic, can become unstratified or homogeneous as a result of an obliteration of the bedding. (2) “False-body thixotropy” in systems where clay minerals and electrolytes are present, results in a “rigid” gel persisting until the yield point is reached, after which “liquid flow” occurs. (3)“Plasticity”, where the water content is within certain limits, causes the material to be a firm gel until the yield point is reached; thereafter, the sediment will fold, shear, microfault and, ultimately, when the water content is sufficiently reduced, it will brecciate. (4) Induration (i.e., cementation) at a comparatively early stage if certain material, such as carbonate, has been precipitated interstitially. This results in fracturing, faulting or brecciation when the stress is rapidly applied, or will flow, fold or flex if the stress is slowly applied. In a water-saturated, lithologically fairly uniform stratigraphic sequence, the value of “m”* (Boswell, 1961) decreases steadily as the overburden pressure increases, so that at any particular locality, “m” decreases with increasing age of the deposits. Exceptions, however, do occur. Table 3-XXXV lists some values of “m” from which the “porosity” or void ratios also have been calculated. There is some correlation between the geologic age of the strata and moisture content, with considerable overlap. Table 3-XXXVI illustrates the general porosity decrease with age of the rocks. In coarse-grained sediments and sedimentary rocks, porosity can be measured

* “m” is conventionally expressed as g of,water per 100 g of solid dried to constant weight at 105-110°C.

249

DIAGENESIS OF SANDSTONES AND COMPACTION TABLE 3-XXXV

Variation in moisture content, porosity, and void ratio with age of argillaceous rocks (after Boswell, 1961, table 1 ) m*

(range)

Devonian, Upper and Middle, marls Carboniferous, Middle, Coal Measures, shales Triassic, Keuper Marl Jurassic, Lias clay Jurassic, Oxford Clay Cretaceous, Gault clay Eocene, Lower Red Beds Eocene, London Clay

Porosity (%)

16

31

9-18

13

26

13 18 22 29 19 24.4

2.5-20 14-22 18-24 28-30 12-27 18-33

25-42

very variable

Void-ratio (e) calculated from m

calculated S and M from rn

9-24

Eocene, Upper 15-29 Oligocene, Upper 19-29 Pleistocene, bedded clays 16-48 usually Postglacial and Recent

rn

(mean)

0.44

(M)

0.35

26 33 38 44 34 40

4.8-20.2 (M) 22.5-27.7 (S)

0.35 0.49 0.61 0.78 0.52 0.66

39 40 41-54

29 (MI

40-82

30-50

4.0-14.6

24 (MI

mode 24.5 23 24

0.64 0.66 0.69-1.17 (S)

0.66-4.6

24-170

* rn = g water/l00 g solid; drying temperature = 1O5-11O0C;values of void ratio, e, and porosity were calculated from rn ; the values from Sorby (S)and Moore (M) are included. either directly by standard techniques such as those used by civil and petroleum engineers, or by indirect methods, such as by determining the natural moisture content “m”, as long as the rock sample is saturated with water when collected. In the latter case, however, the mineralogic composition of the sample should be known, because in the case of clayey rocks, the “porosity” values calculated from “m” will have a different connotation as “m” includes water in the lattices of the clays as well as that adsorbed on their surfaces. Drying temperature will also affect the results (see Cebell and Chilingarian, 1972). Susceptibility to thixotropic behavior increases if the system is essentially loosely packed and a large proportion of the particles is of a diameter less than 1p. The thixotropic properties of sand and silt are greatly enhanced by

K.H. WOLF AND G.V. CHILINGARIAN

250 TABLE 3-XXXVI

Relationship between porosity and age of rocks (after Boswell, 1961, table 2) Age

Porosity (%)

Pleistocene Mio-Pliocene Miocene Upper Cretaceous Permian Pennsylvanian Mississippian Devonian

36-41 31 33-40 24-25 15 15-18 10-1 1 11

the presence of clays. The addition of electrolytes may result either in an increase or decrease in thixotropy, depending on the type of electrolyte and its concentration. An addition of a small amount of clay or other colloidal material, which is strongly hydrophilic, would make a dilatant system more thixotropic. Whereas thixotropy and plasticity are related, dilatancy is antipathetic t o thixotropy. A dilatant state is typified by uniform, close packing of the system’s constituents. Certain natural or artificial systems are either thixotropic or dilatant, depending on the make-up of the system, e.g., percentage of water. The rate at which a disturbed thixotropic system sets varies considerably : some, such as suspensions of montmorillonite clays, can become a gel almost instantaneously, whereas others require minutes, hours, or even days. This time factor should be considered in natural geologic systems, because of the different rates of setting of sediments with a variable composition from one unit to the next, after they were disturbed by an earthquake. Both Boswell (1961) and Niggli (1952), among others, considered the base exchange properties of clay minerals in a system as related to its thixotropy. There is a decrease in thixotropy when electrolytes, such as NaC1, KC1, NaOH, and KOH, are added in certain concentrations to montmorillonite plus water systems, whereas the thixotropy of kaolinite and other clays and powdered minerals increase when the same concentrations of electrolytes are used. One of the authors of this chapter (G.V.C.) found that very small (<1%) concentrations of sodium hydroxide increase thixotropy of both montmorillonite and kaolinite muds, probably because of adsorption of hydroxyl ions on clays. High concentrations of electrolytes usually flocculate the system, greatly increase the viscosity, and decrease thixotropy of clay plus water systems. Upon shaking, the water and clay components in these systems may separate from each other. Much research, however, remains to

DIAGENESIS OF SANDSTONES AND COMPACTION

251

be done on the numerous factors that determine thixotropy. For example, some organic liquids greatly increase the thixotropy of fine sediments, whereas others reduce it. Some hydrophilic, natural organic colloids increase thixotropy by plating the mineral particles with colloidal sheaths. The three main factors which control the degree of thixotropy are: (1)electrostatic or ionic attractions; (2) Van der Waals forces of attraction between molecules; and (3) hydrogen bonds or bridges. In the thixotropic state, the energy of attraction between the particles exceeds that of repulsion, whereas when the mass of material is in the dilatant state, the reverse is true. The surface area of solid components in a system also plays an important role in determining the degree of thixotropy. The area increases as the particle size decreases and as the shape of the constituents departs increasingly from the spherical form. I t should be noted, however, that a thixotropic state can be present in a mass of equidimensional and even spherical grains. The surface area decreases as a result of flocculation of muds (i.e., clay plus water systems) by electrolytes, for example. Creation of gels can be explained by the formation of surface-to-edge contacts in clays, because surfaces of clay plates are negatively charged, whereas the edges are positively charged. Upon shaking, the surface-to-edge contacts are broken and the mass becomes fluid again. Niggli (1952) discussed some properties of sediments which determine compressibility. As pointed out previously, not much detailed information is available on the compressibility and related properties of sediments composed of various proportions of clay, silt, and sand. As the following information demonstrates, it is not sufficient in future research to merely add a certain amount of clay to silt- and sand-sized material for experimental purposes. The exact mineralogical composition of the clay-sized components have t o be known before the results of the experiments can be meaningfully interpreted. Both the adsorbed water content and the extent and type of base exchange that occurs, which mainly determine the compressibility of clayey sediments and, therefore, also of clayey sandstones, also must be known. Figure 3-151 shows diagrammatically the relationship between the range of plasticity (expressed as the percentage of water present in the system) and the type and amount of three different clay minerals (kaolinite, Ca-montmorillonite, and Na-montmorillonite) in clay plus quartz mixtures. With an increase in the content of quartz grains, the plasticity decreases. Figure 3-152 presents the compressibility of quartz, muscovite, kaolinite, and bentonite. One should note, however, that the curve for “bentonite” (an altered tuff; Wolf, 1959) is to be accepted only with care, because the properties of montmorillonite minerals, of which bentonite is commonly composed, varies according to the base exchange cation present (i.e., Caversus Na-montmorillonite). The amount of compaction (in percent, at a pressure of 1kg/cm2) increases with increasing pressure (see Table 3-XXXVII).

K.H. WOLF AND G.V. CHILINGARIAN

252

I

QUARTZ CONTENT,?,

U

u

Fig. 3-151.Plasticity range for various mixtures of quartz with: ( I ) kaolinite, ( 2 ) Camontmorillonite, and (3)Na-montmorillonite. Quartz content diminishes (= reduces) the plasticity range. (After Niggli, 1952,fig. 93,p. 269;courtesy Birkhauser, Basel.) Fig. 3-152.Dependence of the degree of compression on the type of pelitic (= clayey) minerals based on experiments of Haefeli and Von Moos, 1938. A, B, and C are >0.002 mm in size. A = quartz; B = muscovite; C = kaolinite; D = bentonite. (After Niggli, 1952, fig. 94,p. 269 ;courtesy Birkhauser, Basel.)

Table 3-XXXVIII presents the amount of water that can be adsorbed on finegrained materials and the time required to do so. Future comparative studies should determine to what extent the values obtained from laboratory work will apply t o newly-accumulated natural sediments in a number of different sedimentary environments. Table 3-XXXIX demonstrates the dependency of plasticity values on the base exchange capacity of clay minerals, whereas in Table 3-XL another variable, namely quartz content, has been considered. In Table 3-XLI, the grain size of the quartz grains was varied, whereas the amount and type of clay mineral remained constant. There is a general decrease in the lithification, roll, and plasticity values with an increase in grain size. The presence of quartz having diameters less than 2 p changes the. plasticity only little in contrast to when diameters of quartz grains are larger than 2p. In instances

253

DIAGENESIS OF SANDSTONES AND COMPACTION TABLE 3-XXXVII

Specific compressibility (in %) of compressibility a t 1 kg/cm2 as a starting point (after Haefeli and Von Moos, 1938, in Niggli, 1952, table 34, p. 268; courtesy of Verlag Birkhauser, Basel) Material

Pressure (kg/cm2) ~~

~

0.5

1

2

4

-1.51

0

2

4.35

8.3

-3.3

0

3.5

7.4

11.7

-2.39

0

2.7

5.5

8.6

-0.37

0

1.1

2.0

3.1

Platy minerals in low proportions in CaCO3-rich delta sand

0

1.41

3.0

3.5

Platy minerals in high proportions in CaCO3poor delta sand

0

2.6

5.5

8.7

Muscovite, pure

< 2p

Kaoline from Zettlitz, Germany + quartz, 2/3 of the mass of kaoline 200-5OOp in diameter + quartz, 1 / 2 of the mass of kaoline 200-500~ in diameter + quartz, 1/3 of the mass of kaoline 20b-800p in diameter

8

where the components used had diameters less than 2p, however, the plasticity values depended on the types of minerals present (Table 3-XLI1,a).Values for liquid and plastic limits and plastic index for naturally-occurring clay samples are given in Table 3-XLII,b. Sawabini et al. (1974) studied the compressibilities of unconsolidated, fine- to medium-grained arkosic sand cores, lf inches in diameter and 3-4 TABLE 3-XXXVIII Adsorption of water under normal conditions (after Endell, in: Niggli, 1952, table 36, p. 271; courtesy of Verlag Birkhauser, Basel) Type of material

Maximum water content in % dry weight

Fine quartz sand-quartz pelite 27-32 Mica pelite 125 90 Kablinite pelite Ca-montmorillonite 300 Na- montmorillonite 700

Required time in seconds 10-2 3 25 1200 2400 36,000

K.H. WOLF AND G.V. CHILINGARIAN

254

TABLE 3-XXXIX Dependence of plasticity range of the same soil (black soil of India, with 60% particles with d < 2p) after base exchange with different cations (after Oakley, in Niggli, 1952, table 37, p. 271; courtesy of Verlag Birkhauser, Basel) Li-clay Na-clay Mg-clay Water adsorption from N/10 chloride solution; water content (index) a t 110" C 46 Plasticity range (index) in % HzO 82

35 60

17 56

Ca-clay K-clay 17 42

15 22

TABLE 3-XL Dependence of plastic property on the quartz content and type of clay minerals (after Niggli, 1952, table 32, p. 268; courtesy of Verlag Birkhauser, Basel) Ratio of quartz to clay mineral

9:l 7 :3 1:l

0:l

Water content (%) of kaolinite consistency limits

Water content (%) of Ca-montmorillonite

lower

upper

plastici- consistency ty (plas- limits tic index) lower upper

15.5 12.8 12.8 35.7

16.3 19.0 21.6 65.3

0.8 6.2 8.8 29.6

17.8 21.9 49.5

52.5 75.3 140.6

Water content (%) of Na-montmorillonite

plastici- consistency ty (plas- limits tic index) lower gpper 34.7 53.4 91.1

18.5 23.5 47.0

47.5 122.3 214.0 475

plasticity (plastic index) 103.8 190.5 428

TABLE 3-XLI

>

Influence of grain size of quartz on the consistency properties of 50% kaoline (30% 2p; 14% 2-5p; 3% 5-10p; 3% < l o p ) + 50% quartz; quartz grain size was varied from sample to sample as given in the table (after Niggli, 1952, p. 268; courtesy of Verlag Birkhauser, Basel) Grain size of the quartz

Liquid limit Plastic limit Plasticity range (index)

<2p

5-lop

20-5Op

100-2OOp

200-500p

60 40.9 19.1

43 28.5 14.5

36.5 21 15.5

36.8 20.7 16.1

35.1 19.1 16.1

255

DIAGENESIS OF SANDSTONES AND COMPACTION TABLE 3-XLII

<

Plasticity range of material with grain size d 21.1 and naturally occurring clays (after Niggli, 1952, tables 35a,b, p. 270; courtesy of Verlag Birkhauser, Basel) Type of material

Plastic limit

Quartz (Dorentrup, Germany) approx. 28 Calcite, grain size < 21.1, artificially prepared 30 55 Muscovite, grain size < 2p, artificially prepared Kaoline (Zettlitz, Germany) 42 Illite (?) or montmorillonite containing kaoline (Sarospatak) 43 Ca-bentonite with montmorillonite, after Endell 50 47 Na-bentonite with montmorillonite, after Endell Kaoline (Union Co., Illinois): Total sample Lop 0.51.1

36.3 37.1 39.3

Liquid limit approx. 28 48 78 84

18 23 42

120 141 475

77 91 428

58.3 64.2 71.6

Na-montmorillonite (Belle Fourche, South Dakota): 97 Total sample

625-700

Attapulgite (Quincy, Florida): Total sample

177.8

116.6

Plastic index

0

32.3 528-603 61.2

j6

inches long. The cores were obtained from the oil-producing sands of Pliocene age from the Los Angeles Basin, California. Direct measurements of the pore fluid pressure and bulk volume change of each sample were made in a hydrostatic compaction apparatus as the pore fluids were expelled. A t a constant overburden pressure of 3000 psi and a temperature of 140"F, the calculated to 3 * l o d 5 psi-', bulk volume compressibilities ranged from 7.4 * whereas the pore volume compressibilities varied from to psi-' in the 0-3000 psi effective pressure range. The effective bulk and pore volume compressibilities were defined by Sawabini et al. (1974) as follows:

and :

K.H. WOLF AND G.V. CHILINGARIAN

256

where C b e is the effective bulk volume compressibility in psi-'; cpe is the effective pore volume compressibility in psi-'; vbi and Vpi are the initial bulk volume and the initial pore volume of the sample, respectively, in cc; p t is the total overburden pressure in psig; p e is the effective pressure, which is equal to the difference between the total overburden pressure and the pore pressure, in psig; and T is the temperature during the test, in F. Some of the results obtained by Sawabini et al. (1974)are presented in Figs. 3-153,3-154,and 3-155.These authors found that compressibility increases with increasing feldspar content. Somerton and El-Shaarani (1974) studied a group of sandstone and a group of siltstone cores obtained from wells drilled in the Imperial Valley, California, at temperatures as high as 200°C and pressures up to 16,000 psi. Bulk compressibilities of all rocks increased with increased temperature, but the effect of temperature was more pronounced at the lower stress revel (up to =5000 psi). Insufficient data were obtained t o evaluate the effect of temperature on the compressibility of liquid-saturated rocks. Somerton and El-Shaarani (1974),however, stated that the effect appears to be about the same as for dry rocks. Maxwell and Verrall (1954)experimentally investigated consolidation of quartz sand to sandstone, simulating deep burial. Compaction was under5 .-

n

'

c

a

-

K

W

I 3 J

9 0.2 o.z

Y

J

t

3

m W

t

1 u

ld4 01

I '

EFFECTIVE P~ESSURE,p i g

lo4

lo*

b

lo3

EFFECTIVE PRESSURE, psig

10'

Fig. 3-153. Relationship between void ratio and effective pressure for unconsolidated, arkosic oil sands. (After Sawabini et al., 1974, fig. 10, p. 136; courtesy SOC. Pet. Eng. AIME.) Fig. 3-154. Relationship between effective pressure and effective bulk volume compressibility for unconsolidated, arkosic oil sands. (After Sawabini et al., 1974, fig. 6, p. 135; courtesy SOC.Pet. Eng. AIME.)

DIAGENESIS OF SANDSTONES AND COMPACTION

0

s 4

257

J DEFORMATION

Fig. 3-155. Relationship between effective pore volume compressibility and effective pressure for unconsolidated, arkosic oil sands. (After Sawabini et al., 1974, fig. 8, p. 136; courtesy Soc. Pet. Eng. AIME.) Fig. 3-156. Shape of load-versus-deformation curve for sands compacted by direct pressure only. (After Smalley, 1963, fig. 1; courtesy Nature, London.)

taken at room temperature and at 1atm pore-fluid pressure. There was a steady decrease in porosity from about 35%at a total overburden pressure of 10,000 lb/inch2 to about 25%at 50,000 lb/inch2. At a pressure of 50,000 lb/inch ', which is equivalent to 45,000 ft of overburden, the rate of decrease of porosity with increasing pressure was found to be zero. Borg and Maxwell (1956) examined the nature of deformed material, without considering the particle size variation caused by the various pressures. The experimentally simulated high overburden pressure caused complete deformation of samples, which were initially incompletely compacted tind uncemented. On the other hand, Smalley (1963) designed experiments to show the nature of the deformation process as obtained in the laboratory (Fig. 3-156) by measuring the grain size alteration during compression (Fig. 3-157 and Table 3-XLIII), rather than measuring the porosity change. The three stages of deformation observed by Smalley (1963) in a piston compaction apparatus are presented in Fig. 3-156, where load is plotted versus deformation: Stage I depicts elastic compression and the beginning of yield, whereas Stage I1 represents the continued breaking of the sand grains and reduction of the pore space by filling of pores with the broken fragments. Stage 111. sets in when the pores have been filled and the sample becomes in effect a solid, incompressible mass. A 2-inch-diameter steel cylin-

K.H. WOLF AND G.V. CHILINGARIAN

258

LOAD,

tOnS

Fig. 3-157. Variation in particle size percentage in sands compacted at different pressures. (After Smalley, 1963, fig. 2 ; courtesy Nature, London.) Numbers on curves are B.S. sieve sizes.

der with a steel piston was employed t o compress 100-g sand samples at various pressures up to 20,000 lb/inch2. The quartz sand was graded prior to experiments so that 95.9 k 2% passed through a No. 16 and was retained by a No. 18 B.S. sieve. The samples were compressed at a chosen pressure and then analyzed again as to their three-component grain-size distribution, as shown in Table 3-XLIII and plotted in Fig. 3-157. The point of inflexion at the load value along the 18-sieve curve, represents the transition from the Stage I deformation to the Stage 11. The amount of material composed of quartz grains passing the No. 100 sieve steadily increases with compaction pressure. The 22-sieve curve flattens out very early, indicating that there was no increase in this fraction with increasing compression and that the percentTABLE 3-XLIII Size distribution of compacted sands (see Fig. 3-157) (after Smalley, 1963, table 1) Compacting load Compacting pres- % on No.18 B.S. % on No.22 B.S. % passing No.100 (tons) sure (lb/in.2) sieve sieve B.S. sieve

0 2 4 6 8 10 14 26 29

0 1,426 2,852 4,278 5,704 7,130 9,982 18,538 20,677

’

95.5 91.0 61.0 45.5 37.0 32.0 28.0 20.5 20.5

i:

2

4.0 f 2 5.5 15.5 18.5 17.5 18.0 16.5 15.5 15.5

0 0.5 2.5 7.5 10.5 13.5 17.5 25.5 27.5

DIAGENESIS OF SANDSTONES AND COMPACTION

259

age remained the same. The curves suggest that the very fine particles resulting from the breaking of the sand grains fill the pores during Stage I1 of the deformation. Also, it seems that 20% is a limiting value of the No. 18-sieve curve (Le., portion of the sample retained by a No. 18 sieve) and that a further increase in pressure will not change the amount of this size of the sand. Measurements after compression indicated that the assumed original porosity of 35% was reduced to 25% by application of a pressure of 20,000 lb/inch '. Additional experiments were carried out to establish whether or not differently-sized sands would cause any variations in the transition pressures for the various deformation stages. In Table 3-XLIV, Smalley (1963) showed the pressures at the Stage I-Stage I1 transition points in sands having different starting grain sizes. There is a steady increase in this pressure with decreasing particle size. As pointed out by Brooks (1966), the bulk densities of rock masses are important for interpretations of gravity anomalies. In geophysical work, the density variations with depth in sedimentary basins have to be known for precise interpretations and extrapolations of the results obtained. Numerous studies on density variations of fine sediments with depth are available (see Parasnis, 1960, for example), but at the time Brooks (1966) undertook his investigation, data on sandstones were less extensive. He referred to Taylor (1950) who offered some information on the relationship between density and porosity changes with increasing burial. Measurements on natural sandstones have to be compared with the results obtained from laboratory compaction tests on sands, so that one can assess the density distribution in ancient sandstone basins. The necessity of such a comparison becomes even critical, because many investigators claim that fracturing and fragmentation of sands, which occur in the laboratory compaction tests, do not occur in nature. It would be particularly valuable to assess the density distributions, mentioned above, in several basins that contain sediments ranging in age from the Recent through the Precambrian. Inasmuch as this range is TABLE 3-XLIV 1-11 transition pressures in compacted sands (see Fig. 3-156) (after Smalley, 1963, table 2)

B.S. sieve number

1-11 transition pressure (psi)

14-16 16-18 22-25 25-30 30-36 44-52

2570 2670 3200 3750 4050 4800

K.H. WOLF AND G.V. CHILINGARIAN

,

POROSITY %

Fig. 3-158. Relationship between porosity and saturation density (i.e., density of a rock when pore space is filled with water) for sands having different grain densities. (After Brooks, 1966, fig. 1, p. 63; courtesy Geol. Mag.)

rarely available in one basin, by investigating several basins the data can then be obtained step-by-step. Brooks collected samples across the rock sequence, as well as along the dip section, and performed compaction tests on disaggregated sandstone (Bunter Sandstone) to determine the effect of applied pressure on density. The information was used to prepare the density versus depth curve and to assess the bulk density variations in a sedimentary pile. A modified method of Parasnis (1960)for measuring densities was used by Brooks. A rapid method for determination of mineral and grain densities involves the use of a simple linear relationship between the saturation density of a rock and its porosity (= volume of pore space/total bulk volume = vp/vb, in percent), for rocks having the same grain density (see series of lines in Fig. 3-158).This relationship is used because most minerals associated with quartz in sandstones have densities significantly different from that of quartz (2.65g/cm3), e.g., 2.71 for calcite, 2.8-2.9 for dolomite, 2.76 for muscovite, 2.57 for orthoclase, 2.61 for albite, and higher than 5.0 g/cm3 for iron oxides. The densities of plagioclases, oligoclase, andesine, and labradorite, are close to that of quartz. A relationship between saturation density and porosity was shown to exist as follows: If ps = saturation density of rock (pore space filled with water) in g/cm3; pd = dry density (pore spaces empty) in g/cm3; pg = grain (or mineral) density in g/cm3; and r$ = porosity in percent = vp/vb x 100, where V, is the pore volume and vb is the bulk volume, then: r$ = (ps - Pd)

x 100

(3-8)

DIAGENESIS OF SANDSTONES AND COMPACTION

261

Dry density Pd, weight of grains/total volume, is equal to: (3-9) Combining eqs. 3-8and 3-9: =

441 -P g ) 100

+P g

(3-10)

For rocks of the same grain density, therefore, pa is directly proportional 6.The slope of the line is (l-pg)/lOOand the intercept on the pa-axisis pe. The family of curves in Fig. 3-158enables one to determine the average grain density of a rock if ps and 4 are known. Brooks used this diagram in his investigations to detect any significant differences in pg between specimens that could account for any observed density trends. Brooks had to make certain assumptions and mentioned that other factors affect density distribution in a sedimentary sequence, e.g., (a) irregular distribution of cement, (b) neomorphism of minerals during weathering at the surface, (c) presence of unconnected pore spaces, and (d) different grain-size distributions (see also Smalley, 1963). The influence of fluids on density has been discussed by McCulloh (1967)and some of the pertaining information is presented near the end of this section. Brooks' (1966)results, obtained from borehole specimens across the bedding planes and covering a depth range of nearly 1000 ft, are graphically presented in Fig. 3-158.Grain densities range from 2.64 to 2.68 g/cm3, with resulting variation of approximately 25% in saturation density. To reduce this variation to 1096, the grains with densities falling outside the range of 2.66-2.66 g/cm3 were not used. The remaining grains were tested for a possible relationship between ps and depth and a value of r = 0.281 was obtained. This indicates a probability of occurrence by chance of greater than 1 in 10,i.e., the correlation is inadequate. Hence, there is no significant increase of grain density (pg)with depth over the range exhibited by the borehole specimens. The samples collected along the dip and their graphical analysis indicated that the sandstones had a range of grain densities ( p e ) similar to that of the borehole specimens, i.e., across the bedding planes. No significant relationship between ps and the depth along the dip was found. Brooks (1966)performed compaction tests to examine further the possibility of significant density changes with depth by using disaggregated dune sandstones samples. The results of a number of compaction tests are presented in Fig. 3-159,in which density is plotted versus the applied pressure. The dry density values were converted into equivalent saturation densities before plotting, for direct comparison with densities on undisturbed samples. to

262

K.H. WOLF AND G.V. CHILINGARIAN

2.30

ESTIMATED DEPTH OF BURlAL,ft (p=2.25g/cm3)

p

5 m . I

law0 I

15,oOO

I

I

I

1

. l

cE

Pl

c

i

t cn z

2.10

W

n

z 0 Ia

a

2 1.90 I-

a

ul

I

0

1

2.0

4.0

I

SD

1

8.0

PRESSURE, tons/in'

Fig. 3-159. Relationship between pressure (or depth of burial) and saturation density for disaggregated Bunter Sandstone. Curves 1-3 = samples deposited through gauze; curve 4 = sample vibrated to minimum porosity. (After Brooks, 1966, fig. 2, p. 65; courtesy Geol.

Msg. 1

Equivalent depth of burial values are shown on the pressure-axis, representing the approximate thickness of overlying sandstones assuming a bulk density of 2.25 g/cm3. In one set of experiments, the sand was deposited by a free fall from a funnel into the cylinder, resulting in a high initial porosity (40--50%). The results of three compaction experiments are shown as curves I, 2, and 3 in Fig. 3-159. After a preliminary sharp increase in density, the density gradient was reduced to less than 0.04 g/cm3/1000 ft during further compaction. This low value may suggest why significant density increases were not apparent in the natural rock samples collected from the 800 and 600 f t of original depth range, in particular when the influence of grain size and sorting is not known. At a pressure equivalent to an overburden load of about 10,000 ft, the saturation density of samples compacted in the laboratory was always substantially less than the mean value of 2.26 g/cm3 for the natural samples collected. Also, an unloading curve (curve I , Fig. 3-159)shows an elastic relaxation of grains, which resulted in a reduction in density of about 9% of the total density increase due to compaction. Consequently, these curves do not represent the natural compaction of natural rock samples during burial. Inasmuch as naturally-deposited sands usually have a much lower porosity than those obtained by free fall in the laboratory, Brooks (1966) assumed that the unsatisfactory etperimental compaction curves may have been due t o unnatural, initial grain-accumulation packing. Hence, the samples, after

DIAGENESIS OF SANDSTONES AND COMPACTION

263

being dropped into the cylinder, were subjected to intense vibration under water to obtain the minimum porosity. The compression curve 4 (Fig. 3-159) shows that the initial density of the sample approached the natural density. After further compaction and a certain amount of crushing, the sample attained the mean natural density at a pressure equivalent to an overburden of 10,000 ft. Thus, curve 4 represents a more plausible curve of compaction for the natural rock specimens, with a field bulk density of 2.26 g/cm3. The original low porosity of the natural specimens were explained to be the result of wind accumulation, because the sediments of this origin have been shown to have always very low porosities. As Brooks pointed out, water-laid sandstones have different accumulation textures and, therefore, obey different compaction laws, especially if their original porosity was significantly higher than that of the natural sandstone samples used above. Compaction curves I, 2, and 3 (Fig. 3-159) could apply t o some actual water-laid sandstones with sufficiently low densities. These curves indicate that burial to 10,000 f t would result in deformation that would cause a 10% increase in density (not considering long-time effects of compression on pressure solution, for example). On the other hand, density gradients below 1000 f t are very low and could easily be obscured, over limited depth ranges, by other factors not discussed by Brooks. From his tests and theoretical considerations, Brooks concluded that his results are useful in indicating that the averaged densities of surface rock samples are likely to give close approximation to the bulk densities at depth, as long as no unexpected changes in lithology take place. McCulloh (1967), in his important contribution on the volumetric mass properties of sedimentary rocks in situ and the gravimetric effects of petroleum and natural gas reservoirs, has assembled data that is useful for geological and stratigraphic analyses of rock characteristics and in the field of petroleum exploration. Particularly interesting is the fact that he has given consideration to the properties of the various possible subsurface fluids that can alter some of the physical rock properties. McCulloh investigated variations in densities and porosities with depth, and stated that the first-order factors controlling density are mineralogic composition, porosity, composition of the pore-filling fluid, temperature, and fluid pressure. Second-order factors that determine the densities of sedimentary rocks are mainly those that influence the total porosity, e.g., grain size, sorting, depositional environment, depth of burial, postdepositional cementation and recrystallization, deformational history, pore-fluid pressure history, and geologic age. In a Paleozoic section, a few thousand feet thick, composed of consolidated and compacted sandstones, shales, and limestones, an intimate relationship between the dry-bulk density, total porosity, and density profiles over a vertical depth was determined by McCulloh (1967). In a rock with a

K.H. WOLF AND G.V. CHILINGARIAN

264

particular constant grain density, porosity is a straight-line function of drybulk density. The latter ranges from a density of the grains when the porosity is zero to a density approaching zero at 100% porosity. The variations of porosity and density of sedimentary rocks are a function of depth of burial, because gravitational consolidation reduces porosity and simultaneously increases the density of the rocks. These changes are mainly the function of maximum net overburden pressure, degree of compaction, and geologic time. Numerous second-order effects mentioned above, however, also play a role. The interplay of these complex relationships is sufficient to result in extensive variations in porosity versus depth curves from basin to basin and from locality to locality. Such is the case, for example, in Fig. 3-160(Dickey, 1972), where a number of curves for different sandstone units of varying geologic age indicate diverse bulk density and porosity versus depth relationships. McCulloh (1967)obtained a curve of maximum porosity versus depth as shown in Fig. 3-161,by plotting total porosity versus depth for a large number of sedimentary rock samples from a wide variety of geologic environments, represented by a broad range in lithologies of all geologic ages, and from an extreme range of depths. As more data is being accumulated, this curve will probably undergo revision in the future. The greatest variation in porosity is to be expected among the youngest rocks and, especially, newly-accumulated sediments. There is a tendency for later (diagenetic and epigenetic) processes to reduce or even eliminate the minor primary differences in porosity of otherwise litholpgically similar deDEPTH, m

“

T x x ) 2 m 3 o O 0 4 o O 0 ~

0 1

1

0

1

I

1

’

1

UMO

I

I

1

1

8ooo DEPTH, ft

I

“

I

I

‘

12.OOo

I

I

I

0

1 4 m

Fig. 3-160. Relationship between depth o burial and density (as well as porosity) of sediments from various localities. (After Dickey, 1972, fig. 5, p. 7; courtesy Int. Geol. Congr., Montreal.) 1 = Athy (1930), Pennsylvanian, Oklahoma; 2 = Dallmus (1958), Tertiary, Venezuela; 3 = Eaton (1969), Tertiary, California;4 = Dickinson (1953), Tertiary, Gulf Coast;5 = Boatman, Louisiana; 6 = Griffin and Bazer (1969); 7 = Rochon (1967), Louisiana; 8 = Reynolds; 9 = Hedberg (1926, 1936), Tertiary, Venezuela; 10 = Philipp et al. (1963), Germany; 11 = Atwater and Miller (1965), average sandstone, Louisiana.

0

TOTAL POROSITY. IN PERCENT

0

I

jl If

jf iI il

--Probabla maxlmum averaga porosltv of nr arvolr sandstone

if

iI il

probable I---Maximum rosity for most sedC mantary rocks I1 ;I

il

’

if I ;I ;I

-Probabia upper lltnlt of porosity for virtually all sedlmantary rock8

il

il

a of porosities of f I , A a n gmost redlments ynd

-$I I

sedimentary rocks

Fig. 3-161.Relationship between the total porosity of sedimentary rocks and depth of burial, based on empirical data from laboratory measurements of more than 4000 samples of conventional cores from (1)the Los Angeles and Ventura basins of California, (2) other scattered localities in the United States, and (3) the Po Basin of Italy. (After McCulloh, 1967, fig. 4, p. A9;courtesy U.S.Geol. Sum.)

266

K.H. WOLF AND G.V. CHILINGARIAN

posits. The finest-grained and best-sorted sediments tend to be the most porous and least dense upon accumulation, and may contain over 90% interstitial water. In the subsurface, under a few hundred feet of overburden, porosity varies over a wide range, and the coarse-grained and best-sorted sediments, e.g., sandstones and conglomerates, commonly are the most porous and the least dense. At depth beyond a few hundred feet, associated rock units of different lithologies have distinctly different porosities and densities just as they do upon accumulation. Based on extensive research work done recently by numerous investigators (e.g., Roberts, 1969;Chilingarian et al., 1973) one cannot state that clays are more compressible than sands (see Chapter 2, Vol. I). According to McCulloh’s data, the porosity may vary from about 10% at a depth of 20,000 f t to about 30%, and 0.0

DENSITY, IN GRAMS PER CUBIC CENTIMETf 0.5 -.

f t per bbl

cu f t per bbl

I

0-100.000 porn NaCl

I I I

I

I J ,

,

,

,

,

,

,

, I

Fig. 3-162. Variations in density of interstitial fluids as functions of depth, three sets of assumed temperature and pressure gradients, and fluid composition. Curves for interstitial waters from fig. 5 in original paper. (After McCulloh, 1967, fig. 6, p. A12; courtesy U.S. Geol. Surv.)

DIAGENESIS OF SANDSTONES AND COMPACTION

267

occasionally more, at 5000 ft. The greatest ranges in porosity occur in the Pliocene and Quaternary rocks, i.e., the youngest deposits, whereas the porosity of older rocks tend to vary less at all depths. The limits of the range of average porosities in spatially associated clastic marine rocks (i.e., sandstones and argillaceous rocks) with maximum porosities, are given in Fig. 3-161. As a result of variations in mineralogy, the dry-bulk density of sedimentary rocks with zero porosity varies from 2.65 g/cm3 to 2.87 g/cm3. As porosity increases, the variation in grain density will have a decreasing effect

PmbabI. marlmum averam

ecrasltY at l.m n

I-

630

eL

im B

-1

5 10

2 1.6 1.7 Id 1.9 2.0 2.1 2.2 2.3 2.4 2.5 2.6 2.7 FLUIWSATURATED ROCK OENSITY, IN QRAMS PER c u n i c CENTIMETER

0 1.S 1.7 1.8 1.9 2.0 2.1 2.2 2.3 2.4 2.5 2.6 2.7 FLUIO-SATURATED ROCK OENSITY. I N GRAMS PER CUBIC CENTIMETER

8. DEPTH. APPROXIMATELY 4,000 FEET; TEMPERATURE, 1ZO.F; PRESSURE. 1.795 POUNDS PER SQUARf INCH

A. DEPTH. APPROXIMATELY 1.000 FEET: TEMPERATURE. 81’F; PRESSURE,460 POUNDS PER.SQUARE INCH

I-

530

f

I” 2

2r

10

01.6

1.7

1.1

1.9

2.0

2.1

2.2

2.3

2.4

2.5

2.6

2.7

FLUIWSATURATED ROCK DENSITY. I N GRAMS PER CUBIC CEhTlMETER

C. DEPTH. APPROXIMATELY 6.000 FEET: TEMPERATURE. 150’F; PRESSURE. 2.M5 POUNDS PER SQUARE INCH

1.6

1.7

1.8

1.9

2.0

2.1

2.2

2.3

2.4

2.5

2.6

2.7

FLUID-SATURATED ROCK DENSIlY. IN GRAMS PER CUBIC CENTIMETER

D. DEPTH, APPROXIMATELY 10,oOO FEET; TEMPERATURE, 209.F; PRESSURE, 4,465 POUNDS PER SQUARE INCH

Fig. 3-163. Rock density in situ as a function of total porosity and fluid composition at various temperatures and pressures assumed to be prevalent in young deep marine sedimentary basins. Assttmed grain density = 2.67 g/cm3. (After McCulloh, 1967, fig. 8, p. A17; courtesy U.S.Geol. Sum.)

K.H. WOLF AND G.V. CHILINGARIAN

268

on the dry-bulk density, so that two rocks with a 40% porosity and differing in grain density by 0.1 g/cm3, will differ in dry-bulk density by only 0.06 g/cm If these rocks are saturated with the densest of all possible interstitial fluids (namely, a brine), the water-saturated, subsurface densities would also differ by only 0.06 g/cm3. If a rock with 40% porosity is filled with petroleum, instead of salt water, there will be a decrease of 0.2-0.4 g/cm3 in the in-situ bulk density. Consequently, inasmuch as the underground densities of natural interstitial fluids vary widely, their densities must be considered in calculating rock bulk densities. McCulloh (1967)has discussed this aspect at some length, and some of his data in the form of curves are presented here (Figs. 3-162,3-163,3-164,3-165and 3-166). Assuming fixed dry-bulk and grain densities, the density of a rock with a certain porosity is a function only of the density of the fluid filling the pores. Inasmuch as grain densities vary with lithology and fluid densities vary with composition, temperature and pressure, a simple straight-line relationship between porosity and rock density cannot be expected. If one assumes, however, that (a) the porous sandstone is composed of grains having the same density, and (b) the temperature, pressure, and composition of the fluid are constant, then a straight-line graph of rock density versus porosity can be obtained (Fig. 3-163A for a sandstone with grain density of 2.67 g/cm3). The data in Fig. 3-163is applicable to a hypothetical deep marine sedimentary basin having a temperature gradient of 1"F/67f t and a pressure

'.

F

40 Ic W z

35

0

a 30 W

a z 25 W a $20

88 2

W

& \

52

6000

*<

8500

*%

15

< 10

I- -

20,000

7 -

?

:!5 0

1.7 1.8 1.9 2.0 2.1 2.2 2.3 2.4 2.5 2.6 2.7 FLUID-SATURATED SANDSTONE DENSITY. I N GRAMS PER CUBIC CENTIMETER

1.6

Fig. 3-164. Relationship among average density in situ of reservoir sandstone, its fluid composition, and total porosity. Assumptions: a. probable maximum average porosity for depth, temperature, and pressure gradients prevalent in young marine sedimentary basins; b. gas-oil ratios as given in Figs. 3-162 and 3-163. (After McCulloh, 1967, fig. 9, p. A18; courtesy U.S.Geol. Sum.)

DIAGENESIS OF SANDSTONES AND COMPACTION

269

DENSIM, IN G R A M S PER CUBIC CENTIMETER

0

POROSITY, IN PERCEM

5

0

0

EXPLANATION

Densityfidd br many cornrnercially important ptmleurn fluids

Density field for most interstitial waters

Density field of sedirnentory Density field d sedimentary rocks saturated with wrocks saturated with troleum fluids woter

POCOSity fieU of sediMntS and sedimentary rocks

Fig. 3-165. Densities and porosities of sedimentary rocks and their constituents as functions of depth and of temperature and pressure gradients prevalent in young deep marine sedimentary basins. (After McCulloh, 1967, fig. 10, p. A19;courtesy U.S.Geol. Sum.)

gradient of 0.445 psi/ft. Similar density-depth curves can be calculated for other temperature and pressure gradients. In Fig. 3-163A,the sandstone is assumed to be saturated at different times with each of the four principal interstitial fluids. The densities of these fluids are plotted versus depth in Fig. 3-162,under temperature and pressure conditions common at a depth of 1000 f t in young and deep marine sedimentary basins (indicated as “normal” in Fig. 3-162).The four lines in Fig. 3-164 converge on the grain density at zero percent porosity, as the lower porosity results in the reduction of density contrast between rocks saturated with water and petroleum. According to Fig. 3-161,the maximum sandstone porosity in young’sedimentary basins is 50% at a depth of 1000 ft, whereas the average porosity of these sandstones at shallow depth is about 40% or

K.H. WOLF AND G.V. CHILINGARIAN

270

5 ;

40

0

35

'

w' 30

z

n

-2

a

v)

25 20

LL

15 l-

a

a 0 10 0 " 5

-1

2O

n

+

-0

-0.1 -0.2 -0.3 -0.4 DENSITY CONTRAST, IN GRAMS PER CUBIC CENTIMETER

Fig. 3-166. Density contrasts between reservoir rocks saturated with water and those saturated with petroleum fluids, as functions of hydrocarbon composition and average total porosity, assuming temperature and pressure gradients prevalent in deep young marine sedimentary basins. (After McCulloh, 1967, fig. 11, p. A20; courtesy U.S.Geol. Surv.)

less. The horizontal line at a 40% porosity intersects the sloping straight line of methane-saturated rock at the lowest density of 1.61 g/cm3 (Fig. 3-163A). This suggests that the rock density lower than 1.6 g/cm3 at a depth of 1000 f t is probably due t o the presence of coal, diatomites, or rocks of unusual composition. Figure 3-163B is analogous to Fig. 3-163A, except that the temperature and pressure are those common at a depth of 4000 f t , according t o Fig. 3-162.According to Fig. 3-3.61, the maximum sandstone porosity at a depth of 4000 f t is less than 40% and averages about 30%. If a rock at that depth has a density less than 1.90 g/cm3, it must be due to the presence of coal, a rock of unusual composition, or an overpressured formation (a rock dilated by abnormal fluid pressure). The graphs in Figs. 3-163C and 3-163D supply similar data for depths of 6000 and 10,000 ft, respectively. Relationship between the average sandstone density and depth is presented in Fig. 3-164. The temperature, pressure, and porosity gradients used in plotting this figure are nearly limiting values, i.e., temperature and pressure at any depth are near the minimum values found in boreholes, whereas porosity at any depth is near the maximum. Although an increase in temperature would result in a less dense fluid, it would probably be accompanied by a decreased porosity, thus compensating for each other. Figure 3-165 summarizes the data assembled by McCulloh (1967) on the relationship between the depth of burial and densities of sedimentary rocks.

DIAGENESIS OF SANDSTONES AND COMPACTION

271

Fig. 3-167. Density contrasts between water-saturated argillaceous rocks of various maximum porosities and reservoir sandstones saturated with water or petroleum fluids, as functions of average total sandstone porosity and fluid composition. (After McCulloh, 1967, fig. 12, p. A21; courtesy U.S.Geol. Sum.)

In Fig. 3-166,the density data presented in Fig. 3-164have been replotted in the form of density contrasts. Water-saturated porous reservoir rocks are used as a standard against which the densities of same rocks saturated with various fluid hydrocarbons are contrasted. The “total” porosity is the average porosity and, as in Fig. 3-164,the values of average total porosity have been correlated with probable maximum depths by reference to the porosity data in Fig. 3-161.Figure 3-167demonstrates that the greatest density contrasts between water-saturated sandstones and water-saturated argillaceous rocks occur where the sandstone porosity is maximum, i.e., in the youngest, uppermost rocks in a stratigraphic section. EFFECTS OF COMPACTION FLUIDS ON CLAY MINERALOGY IN SANDSTONES

It is now an accepted fact, supported by laboratory and petrologic observations on natural rocks, that clay minerals within sandstones (e.g., graywackes, kaolinitic quartzites, montmorillonitic pyroclastics or volcanic arenites, and argillaceous arkoses) can be the product of: (a) detrital origin, with or without secondary alteration in situ; (b) direct precipitation from fluids within the sandstone framework; (c) neomorphic replacement of unstable minerals, such as feldspar; and (d) any combination of the above. Fluids, like compaction solutions, are most important in the processes in-

272

K.H. WOLF AND G.V. CHILINGARIAN

volving clays in sandstones, because they supply and remove chemical components and make the environment conducive for reactions requiring a certain pH and Eh (Dapples, 1967). In other sections of this chapter, several publications were discussed that present data on the release of fluids from the clays during diagenetic mineralogical transformations (e.g., Powers, 1967; Fyfe, 1973). Some of the data concerning the influence of fluids on the composition of clays, which is a somewhat different, though related, topic from the release of water by clays during diagenesis, are presented here. Hiltabrand et al. (1973)performed diagenesis experiments on argillaceous sediments in artificial sea water and recorded mineralogic and chemical elemental changes with variation in temperature, depth, and geologic age. They concluded that the mineralogy of mudstones and shales is the result of diagenesis rather than a reflection of the chemical milieu of the depositional environment. This conclusion also applies to clays present as interstitial matrix within sandstones, as long as the clays are detrital in origin and diagenesis was related to the near-surface conditions of the depositional milieu. It may not apply to neomorphic clays because they can originate much later through the interaction of unstable elastic grains and compaction fluids, for example. If ion exchange is one of the processes involved, then the reader is referred to the publication by Carroll (1959)who described ion exchange of all major clay mineral groups, some other minerals (chlorite, glauconite, allophane, feldspar, quartz, and zeolites), and some rocks, e.g., basalt, tuff, shale, and bentonite. Sarkisyan (1971)used a scanning microscope in the study of clay cements in elastics of an oil-bearing basin to determine reservoir properties. The reservoir rocks are present at different depths and form zones of postsedimentary (= epigenetic) alterations of rocks so that zones of kaolinite, kaolinite-chlorite, and chlorite cements were distinguished. The properties of the allogenic (= detrital) and authigenic (= neoformed) clay cements as determined by electron-microscope examination allowed differentiation between the two. Towe (1962)discussed the derivation of silica from clay deposits as a source of cement in sandstones. He stated that with an increase in depth of burial and under certain geochemical conditions, Si4+ can be expelled from the tetrahedral layers as it is being replaced by A13+ in the montmorillonitetype clays and thus cause their alteration to micas. This would result in an increase in illite content while silicon will be released during this transformation process. Inasmuch as it has been shown that both the depth of burial and the geothermal gradient appear to have an effect on such changes, it seems quite plausible to assume that compaction and compaction fluids are involved. Towe, using published data, showed that 3 g of silicon are released per 100 g

DIAGENESIS OF SANDSTONES AND COMPACTION

273

of pure clay in the transformation of a montmorillonite (bentonite) to an average illite, wheras 1 g/100 g is freed in a conversion of mixed-layer illite-montmorillonite to illite. Hence, there is a sufficient amount of silica available to account for the cement in sandstones, but the amount of silica available varies depending on several geologic conditions mentioned by Towe. One of the main controls would be the type of clay mineral supplied from the terrigenous source-rock area. If the latter had furnished montmorillonite to a marine environment, the resultant clay deposit should be a good source of silica, whereas if illite clay had been supplied, only a poor silica source would be available in the depositional environment. Whetten and Hiltabrand (1972) used recent, coarse, silt-free and clay-free sand from the Columbia River, composed largely of andesitic volcanic detritus, and performed a hydrothermal experiment at a temperature of 200°C and a pressure of 200 psi in a brine solution. After five months, the clastic grains were altered to a mixture of 82% (by weight) sand, 5% silt, and 13% montmorillonite clay. Although only the hypersthene grains were severely etched, other mineral grains and lithic fragments probably also participated in the reaction to form clay. The clay, which coated various parts of the apparatus, appeared to have at least partly precipitated from solution. The ' , a slight decrease in Ca fluid phase showed a slight increase in Na+ and K and Mg2+, and a great increase in Si4' content (saturation); pH remained constant at 3.5. These results suggest that the matrix of graywacke sandstones could, at least partly, be of secondary origin, as has been proposed by Cummins (1962) and is shown in Fig. 3-168. This supposition is also supported by the geochemical investigation of Reimer (1972), who found that the matrix in Precambrian graywackes formed through reassemblage of

'+

Stable sand grain

Rather unstable sand grain

Very unstable sand grain

Interstitial water

Fig. 3-168. Diagram showing presumed post-depositional origin of matrix of graywackes. (After C u m i n s , 1963, fig. 4; in: Pettijohn et al., fig. 6-20, p. 210; courtesy Springer, New York.)

K.H. WOLF AND G.V. CHILINGARIAN

274

chemical compounds derived from decayed labile clastic constituents such as feldspars, mafic minerals, and fragments of igneous and mafic volcanic rocks. The resulting matrix consists of chlorite, sericite, and dolomite in varying amounts. When the matrix plus dolomite content is plotted versus the content of labiles, a volume increase of 1.6 to 1.9 parts of matrix plus dolomite for 1 part of labiles is indicated. This suggests an outside source, and transportation of the constituents by pore fluids into the graywacke sandstone. The Ca2+ cation was partly derived from the decomposition and partial albitization of the plagioclase within the graywackes, whereas the Mg2+ was released from the basic igneous rock grains within the deposits. Most of the Ca2+ of the dolomite, however, was derived from the destruction of the plagioclase in shales alternating with the graywackes and was brought in by pore fluids. This is indicated by the peculiarities in the Rb/Sr ratios of both graywackes and shales. The CaO of the average shale was reduced from about 1.64% to 0.19% and the Sr content from about 81 ppm to 25 ppm, whereas at the same time the CaO content of the graywackes increased from about 1.64% to 1.89% and the Sr content from 81 to 91 ppm. The COz gas was almost totally introduced by pore fluids. Bucke and Mankin (1971;see also Wilson, 1971) studied the illite, chlorite, and kaolinite present in both sandstones and shales of mid-Pennsylvanian age and found consistent differences between the types and contents of clays in the shales and sandstones (Figs. 3-169,3-170,and 3-171).The shales Illite

0

Chlorite

Kaol i ni t e

Fig. 3-169. Clay mineral composition of Desmoinesian (Middle Pennsylvanian) sands and shales (finer than 4 p). Open ciicles = sand; solid circles = shale. (After Bucke and Mankin, 1971, fig. 2, p. 974;courtesy J. Sed. Petrol.)

DIAGENESIS OF SANDSTONES AND COMPACTION Sholt S o d

OILLITE ~'-JCHLORITE

275

Shale SDnd

KAOLINITE

Fig. 3-170. Comparison of clay-mineral content of adjacent sand-shale pairs (finer than 4 p). (After Bucke and Mankin, 1971, fig. 3, p. 974; courtesy J. Sed. Petrol.)

have a limited compositional range and illite dominates in all samples. Cornpositional variation is prevalent in sands, but illite is still the major constituent. Bucke and Mankin closely investigated associated sandstones and shales, because they considered this approach to be a prime starting point for establishing the relative importance of detrital versus diagenetic origin of clay

SAND untreated

+SHALES-

+SANDS-

Size in microns

SAND

solvatcd

SHALE untreated SHALE solvoted

0ILLITE

CHLORITE

~KAOLINITE

Fig. 3-171. Variation of Clay-mineral content in Desmoinesian sands and shales with size fraction. (After Bucke and Mankin, 1971, fig. 4, p. 975; courtesy J . Sed. Petrol.)

276

K.H. WOLF AND G.V. CHILINGARIAN

minerals. Various textural and other considerations led to the conclusion that kaolinite is authigenic, whereas illite and chlorite are detrital in origin. It should be noted here that the sandstones are much higher in kaolinite content than shales. Bucke and Mankin (1971,p. 975)stated: “The illite and chlorite were subjected only to some minor ‘repair work’ or reconstruction diagenesis . . . Those clays deposited in shales or mudstones have been preserved with only minor reconstruction by ion absorption and exchange. Any changes that may have occurred were probably physical changes, notably water expulsion during compaction as described by Burst (1969).Those weathered illites and chlorites deposited with coarser material remained in a post-depositional environment with continued relatively high porosity and permeability even after compaction. These clays had prolonged access to the required ions, notably potassium and magnesium, brought by circulating interstitial water thereby reversing some weathering effects.” Using the paragenesis of several authigenic minerals, i.e., quartz, kaolinite, calcite, dolomite, and pyrite-siderite, Bucke and Mankin suggested a trend toward a higher effective pH during diagenesis. It must be pointed out, however, that although Zimmerle (1963)and Sharma (1968,1970) based their interpretation of diagenetic alterations on pH, Bucke and Mankin stated that the actual change in pH may have been minor because an increase in temperature has much the same effect on solubility of quartz and calcite as an increase in pH. Carrigy and Mellon (1964)described authigenic clay mineral cements in sandstones from Alberta. Often no clear distinction can be made between detrital and authigenic clays; indeed no sharp division was possible between grains, cements, and matrices among the fine constituents. Nevertheless, certain useful results were obtained. The stratigraphic section studied is dominantly non-marine with a maximum thickness of about 25,000 f t prior to uplift and erosion. A generalized section is given in Fig. 3-172,which also shows the distribution of the authigenic clay minerals and zeolites. The strata markedly decrease in thickness to the east from the Foothills to the Plains, where they form a succession of nearly flat-lying, interbedded marine and non-marine strata, which are several thousand feet thick and extend north and south to the margin of the underlying Precambrian basement. The details related to the fine-grained minerals are as follows (Carrigy and Mellon, 1964,pp. 468-47 0): (1)Kaolinite is the most widely-distributed clay mineral cement. It is common both in marine and non-marine sandstones and is present in both the deeply buried, folded strata of the Foothills and in the shallow, flat-lying strata of the Plains. Kaolinite is the dnly common authigenic clay in quartzose sandstones low in feldspathic and volcanic detritus, where it is associated

DIAGENESIS OF SANDSTONES AND COMPACTION

277

LEGEND

non%&ne

and shale

morine sandstone

-

marine siltstone and shale unconformity

z - zeolites c -chlorite

k - koolinite i - illite m - montmorillonite

Fig. 3-172. Generalized columnar section of the Cretaceous-Tertiary strata of the Rocky Mountains and Plains regions of Alberta showing the known distribution of authigenic clay mineral and zeolite cements in sandstones. (After Carrigy and Mellon, 1964, fig. 4, p. 469; courtesy J. Sed. Petrol.)

with authigenic quartz and, generally, calcite. Hence, kaolinite in quartzose rocks originated from the chemical precipitation from the alumina-rich compaction solutions expelled from the adjacent shaly strata and is not the alteration product of feldspars. Kaolinite is also abundant in many sandstones with moderate amounts of volcanic detritus, where it is found with other cements, excluding zeolites. The chlorite-kaolinite assemblage is common in the fluviatile sandstones, but only kaolinite is abundant in sandstones interbedded with coal or marine shales in these units. Kaolinite is also present as the main clay cement in sandstones associated with marine and non-marine sandstones. Montmorillonite and glauconite are also common in many of the above sandstones; however, both minerals, especially glauconite, appear to be associated largely with altered clastic grains (probably pyroclastics?) in contrast to the intergranular occurrence of the authigenic kaolinite. (2) Authigenic illite is less widely distributed than kaolinite. It was found

K.H. WOLF AND G.V. CHILINGARIAN

278

only in the Foothills as interstitial cement in some non-marine fluviatile sandstones, which contain moderate t o abundant amounts of feldspar and volcanic clasts. In these sandstones, illite is commonly associated with other authigenic components: chlorite or kaolinite, quartz, and calcite. Illite is also present in one basal, marine, feldspathic sandstone member, cemented by a mixture of kaolinite and fibrous illite in combination with quartz and calcite. (3) Authigenic chlorite is abundant in many volcanic sandstones and is associated with other authigenic minerals: laumontite, illite, kaolinite, montmorillonite, quartz, and calcite. Chlorite is scarce or absent in coal-bearing or marine facies. It is also scarce or absent in the more deeply buried sandstones many of which contain abundant volcanic clastics, but which have kaolinite and/or montmorillonite as cements. Relation between types of clay minerals and depth of burial is probable and should be investigated. (4)Montmorillonite is present in many marine and non-marine sandstones; however, its distribution as an intergranular cement, in contrast to its presence as an alteration product within clastic grains, is limited. Where present, montmorillonite is invariably found in volcanic-rich sandstones, although these sandstones do not necessarily contain montmorillonite. The suggested origin of the four groups of authigenic minerals described above and correlations between the gross composition and inferred depositional environments, are presented in Fig. 3-173. Assemblages of authigenic silicates found in the shallow, undisturbed sandstones of the Plains are less diverse than those observed in the folded sandstones of the Foothills. In the opinion of the present writers, these differences in regional distribution of the authigenic clay cements have occasionally been attributed to differences in the physicochemical conditions and depth of burial to which the rocks were subjected, rather than to differences

MARINE AND NE AR-MARINE

illite

- kaolinite - q u a r t z

kaolinite - q u a r t z

kooliniie

-94

V O L C A N I C ROCK FRAGMENTS AND FELDSPARS-

F e l d s p a t h l c greywackes

P r o t o q u a r t z i tes

Fig. 3-17 3. Chart showing the relationship among authigenic silicate assemblages, bulk composition, and depositional environment of Cretaceous-Tertiary sandstones of the Alberta Foothills. (After Carrigy and Mellon, 1964, fig. 5, p. 471; courtesy J. Sed. Petrol.)

DIAGENESIS OF SANDSTONES AND COMPACTION

279

in the original composition and sedimentary environment. This assumption seems incorrect, however, because Carrigy and Mellon stated (p. 471): ". . . within the Foothills proper, where the sediments have been subjected to geosynclinal depths of burial and subsequent tectonism and, by inference, to widely varying conditions of pressure and temperature, authigenic silicate assemblages show no obvious correlation with original depths of burial or intensity of folding. Instead, they are intercalated throughout the stratigraphic succession in relation to sandstone composition and, to a lesser extent, the inferred depositional environments of the rocks, demonstrating the stability of all four major types of clay minerals over the range of temperatures and pressures prevailing during diagenesis. Presumably, the ultimate factor responsible for their formation was the composition of the pore fluids during cementation, as determined by the availability of the component chemical constituents from the surrounding detritus." Meade (1963, 1964) found an anomalous increase in pore space with increasing depth when one would expect a decrease in porosity. Grain-size variation could not explain this anomaly, so that Meade searched for additional factors through a literature survey and examination of cores. He concluded that the following variables, in addition to overburden pressure, have had an influence on pore volume in fine sediments: particle size, clay minerals and their adsorbed cations, concentration of interstitial electrolyte solutions and their acidity, clay-particle orientation, and presence or absence of microfossils (i.e., diatoms in this case). Figure 3-174 summarizes some of the factors of interest here in relationship to compaction and other diagenetic processes. Figures 3-174, a and b demonstrate the influence of particle size on poor volume: in both unconsolidated marine sediments, and sediments under overburden pressure, the pore volume increases with decreasing particle size. Figure 3-174,c shows that in the pressure of 1-100 kg/cm2, the sequence of increasing pore volume is kaolinite + illite + montmorillonite (see also Chilingar and Knight, 1960). This is probably the result of particle size, 'because the specific surface measurements show a decreasing order of particle size in the same sequence. Figures 3-174,d, e, and f are of particular interest to illustrate a possible relationship between chemical diagenesis and overburden pressure (i.e., compaction). In sandstones with clay minerals, either as a matrix or as lenses or beds, the exchangeable cations adsorbed by clays are also influencing pore volume under low-to-moderate burial pressures as shown in Fig. 3-174,d. Meade (1963) mentioned that the smaller the valence and the larger the hydration radius of the adsorbed cation, the greater the pore volume of the clay sediment. This does not seem to be applicable under large overburden loads (30-3200 kg/cm '), as demonstrated by experiments performed by

K.H. WOLF AND G.V. CHILINGARIAN

280 a

h

O

0 - O

C

10

Average partide size (9)

Kaolinite

l 100

10

Pressure

Effective overburden load ( kg/cm2)

(kg/cm2)

e

d

f

47

21

> 0 10

100

Pressure t kg/crn2 )

0

1 1

100

l b M , NaCl

,

10

100

Pressure (kg/cm2)

K 10-'M NaCl

0

10

1 100

Pressure !kg/cm2)

Fig. 3-174.Relations of void ratio to other factors, observed in natural sediments and in laboratory experiments. Void ratio is ordinate in all graphs; note different void-ratio scales. a. Relation to average particle size observed in unconsolidated sea-bottom sediments. Curve I modified after Von Engelhardt (1960,p. 15); curve I1 modified after Shumway (1960,p. 663). b. Generalized relation to effective overburden load and particle size in sediments. Modified after Skempton (1953,p. 55). c. Experimentally determined relation to pressure and clay-mineral species. Modified after Chilingar and Knight (1960,p. 104), to show their results to 100 kg/cm2. d. Experimentally determined relation to pressure and adsorbed cations in <0.2 p fraction of montmorillonite. Modified after Bolt (1956,p. 91). e. Experimentally determined relation to pressure and electrolyte concentration in unfractionated Fithian illite (about 60%by weight coarser than 2 p ) . Modified after Mitchell (1960,fig. M3). f. Experimentally determined relation to pressure and electrolyte concentration in < 0.2p fraction of Fithian illite. Modified after Bolt (1956,p. 92). (After Meade, 1963,fig. 2, p. 237;courtesy Sedimentology. )

Von Engelhardt and Gaida (1963). Figures 3-174,e and f illustrate the interrelationships among overburden pressure, degree of compaction, particle size, and electrolyte concentration. Further research work is needed in this area.

DIAGENESIS OF SANDSTONES AND COMPACTION

281

EFFECT OF COMPACTION FLUIDS ON TRACE-ELEMENT AND ISOTOPE COMPOSITION OF SEDIMENTS

The control of compaction fluids on trace-element and isotope composition of sediments is a highly diversified field of investigation. It can only be treated briefly here by considering a few specific examples and by some discussions related to the numerous complexities involved in genetic interpretations. The trace-element budget in a sandstone, that is related to the intergranular clay matrix and can, therefore, be influenced by compaction (or any other) fluids, is considered first.

The trace-element budget It has been shown by numerous investigators that: (1)detrital clays can undergo diagenetic ion exchange at the site of sedimentation; (2) detrital clays can be diagenetically altered partly to completely to new (neomorphic) clay minerals and release certain ions (this process may continue into the metamorphic stage); (3) labile clastic grains can alter into clay minerals during diagenesis and burial metamorphism; and (4) neomorphic clay minerals can be the result of chemical precipitation from solution, both during diagenetic and metamorphic stages. Any of the above clays may undergo more than one generation of changes in elemental composition, depending on various physical and chemical factors changing through geologic time. Compaction fluids may cause ion exchange and bring metallic ions that were not available earlier, or they may remove elements that were released into the fluid. In the latter case, diminution or depletion of ions would occur. The composition of the clays in cases 2 to 4 can be influenced by the compaction solutions. The boron content Keeping in mind the above discussion, the boron content of sediments is considered next. One could have selected a number of other elements, but the geochemical behavior of boron is particularly important because it is used as paleosalinity and paleotemperature indicator. Harder (1970) stated that the geochemical distribution of elements in sedimentary rocks is genetically complex as they occur in the following forms: (1)in clastic fractions in the sand, silt or clay; (2) adsorbed on fine particles, e.g., clays; (3) chemically precipitated fractions as newly-formed minerals; and

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K.H. WOLF AND G.V. CHILINGARIAN

TABLE 3-XLV Boron content of oxides and silicates (after Harder, 1959, in Harder, 1970, table V, p. 162) Mineral

Boron content (ppm)

Remarks

Quartz Agate

0-35 2-90

Nesosilicates Cyclosilicates

several ppm, sometimes under the detection limit

Phyllosilicates

distinctly higher than in the other see Table 3-XLVI minerals, especially the pure, clean micas, serpentine and montmorillonite

Tectosilicates

several ppm, the higher values in sericitized plagioclases and in minerals of the sodalite and scapolite groups

approximately, mean value distinctly higher than quartz

(4) organically precipitated fractions. Boron is found in all four cases. The boron content in different sediments, different grain-size fractions, and minerals is presented in Tables 3-XLV, 3-XLVI, and 3-XLVII. Illites contain the most boron, kaolinites the least, and montmorillonites and, probably, chlorites have intermediate contents TABLE 3-XLVI Boron content of phyllosilicates (after Harder, 1959, in Harder, 1970, table VI, p. 162) Mineral

Boron content (ppm)

Muscovite Paragonite Biotite Sericite Illite Glauconite Montmorillonite Kaolini te Serpentine Chlorite

10-500, very variable values 50-250 1-6 40-2000 100-2000 or more several thousands

5-200

10-30 low or very high values, according to conditions of formation around 50, but generally lower

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283

TABLE 3-XLVII Boron content of sediments (after Harder, 1970,table I, p. 155) Rock

Clays and shales Sandstones and sands Limestones Dolomites Ironstones Glauconitic rocks Possible average for sedimentary formations

Boron content (ppm) range

mean value

25-800 5-70 2-95 10-70 20-200 350-2000

100 35 27 28 85

(see also Couch, 1971). Boron content is greatly dependent on the type of facies, with clays being the richest in boron. The boron content in limestones is controlled by the amount and type of clays present, whereas in sandstones it depends on the amount of clay and the presence of tourmaline. Especially micas and sericitized plagioclase are rich in boron. Fine-grained iron oxide or hydroxide precipitated in a marine environment can contain up to 300 ppm boron, whereas only a few ppm are present in lacustrine deposits. Chamosite has similar boron contents, but glauconitic rocks are richer in boron. Glauconite, which is a diagenetic neomorphic mineral produced through recrystallization, has a crystal structure that can easily accommodate boron. Its boron content is higher than that in chamosite and iron ores. In iron oxide, iron hydroxide, and iron silicate (except glauconite), the boron is probably adsorbed at the surface. This weak type of fixation is responsible for the sharp decrease in boron content after even surficial weathering. The effects of diagenesis and metamorphism are even more pronounced, as discussed below. The boron content of clay minerals depends on a number of variables: (1) boron content of the fluids, e.g., continental and marine waters, volcanic solutions, and surface and subsurface fluids (Fleet, 1965); (2) time of exposure of clays to the fluids (Fleet, 1965); (3)rate of sedimentation (related to 2 in early stages of diagenesis; Fleet, 1965); (4)salinity (Fleet, 1965; Couch, 1971); (5) water temperature (Fleet, 1965); (6) common-ion effect; (7) specific-surface area (controlled by grain size; Couch, 1971); (8)type of mineral (Couch, 1971); (9) crystallinity (related to 8;Couch, 1971); (10) pH of the interstitial solution (Fleet, 1965); (11)inherited boron or boron on recycled clays, which will retain boron from its earlier history (Fleet, 1965; Perry, 1972; Couch, 1971); (12) organic content of the sediments (Fleet,

284

K.H. WOLF AND G.V. CHILINGARIAN

1965; Walker, 1964). (The above selected references are among the many publications available on these topics.) Although many investigators discussed the conditions close to the surface of sedimentation that control the boron content in clays, many of the same factors should be influential in the subsurface during higher-grade diagenesis and metamorphism. The following factors control the release or take-up of boron by clays: (1)time of exposure available (depending on rate and degree of compaction which, in turn, influence porosity and permeability); (2) vertical changes in composition of the interstitial and compaction fluids; (3) temperature; (4) pressure; (5) recrystallization; and (6)pressure solution (which may release boron from quartz). Based on experimental results, Harder (1970)showed that the uptake of boron by illite, for instance, occurs in two steps: (1)adsorption of boron on the surface of the clay mineral which proceeds very quickly; and (2) subsequent boron incorporation in the tetrahedral sites of illite structure, which is a very slow reaction. Release of the adsorped boron is easier, therefore, than that of the absorbed boron within the structure. Harder (1961)stated that diagenesis and metamorphism through recrystallization and neomorphism decrease the boron content of sediments. This boron is added to the “hydrosphere” and returns to the sea, thus completing one part of the boron geochemical cycle. Diagenesis generally reduces the initial boron content of sediments with time, because of the increase in grain or crystal size through recrystallization and neomorphism. This has been observed in the case ‘of recent, volcanic, submarine iron-hydroxide precipitates that contain 300 to 400g of boron per ton obtained from the sea water, whereas in Devonian iron ores of a similar genetic type, the boron content is only 18 g B/ton of precipitates. Crystallization and recrystallization of the ferric hydroxide to hematite releases most of the boron. Similar trends of boron decrease have been observed in purely sedimentary, syngenetic oolitic iron ores, as well as in iron silicate minerals. The latter were formed through diagenetic reduction of ferric hydroxide to ferrous hydroxide and subsequent reaction of the Fez+ with aluminium and silica. The geologically young chamosite contains 60 g B/ton of ore, the Devonian ores contain 50 g B/ton, and the partly chloritized ores of Ordovician age have only 13 g B/ton. In the latter case there is a change from the kaolinitetype structure of the chamosite to the chlorite structure. If the iron ores underwent higher-grade metamorphism so that the whole rock has been chloritized and contains garnet, feldspar, actinolite and magnetite, then the boron content is very low (e.g., less than 5 g/ton). Where the boron is diagenetically released into circulating subsurface fluids, it may form authigenic tourmaline in unmetamorphosed sediments. Perry (1972) studied boron uptake into the lattice of clays during the

DIAGENESIS OF SANDSTONES AND COMPACTION

285

deep-burial diagenesis stage. Such a study seems reasonable, according to him, because experimental results have shown increased boron fixation at higher temperatures (Couch and Grim, 1968) and that many illites are diagenetic in origin. Perry examined samples from deeply buried Gulf Coast sediments and determined that the starting clay mineral assemblage was composed mainly of a highly expandable, mixed-layer illite-montmorillonite with lesser amounts of kaolinite, discrete illite, and sometimes chlorite. With increasing depth of burial, the montmorillonite layers were converted to illite layers and the proportion of illite layers in the mixed-layered illite-montmorillonite increases from 30 to 80% with increasing depth of burial. The boron content of the less than 1-mu fraction from two wells showed only a vague trend of increasing boron concentration with depth. In another approach, Perry plotted the boron concentration in the less than 1-mu fraction versus the proportion of illite layers in the mixed-layer illitemontmorillonite. He found that the relationship is much more clearly defined in this case and that the amount of boron fixation in the fraction less than 1mu in size is best related to the diagenetic formation of new illitic layers. According to E. Couch (personal communication, in: Perry, 1972, p. 156), clay minerals being supplied by the Mississippi River contain 50-80 ppm boron. These concentrations are close to the lower values found by Perry for the highly expandable clays. Through diagenesis, the boron content for the illite-montmorillonite increased to nearly 200 ppm. By extrapolation of this data, a boron value of 270 ppm was obtained for a non-expandable illite, which falls well within the range of boron contents of the marine illites of sedimentary rocks (Reynolds, 1965; Harder, 1970). Perry (1972, p. 156) concluded that ". . .the diagenetic illite formation and its concomitant boron uptake therefore can account adequately for the boron increase necessary to indicate a marine environment. With increased time under diagenetic conditions, boron would continue to diffuse into the illite structure in order to reach higher boron values found in some older illites. The higher temperatures and more concentrated pore fluids encountered in diagenetic conditions provide more reasonable conditions for boron incorporation in the lattice than the conditions at the watersediment interface, and the diagenetic formation of new illitic material is certainly more favorable." Perry (p. 156) continued to state that ". . . pore water flow could increase the amount of possible boron gain from pore waters, but only if the flow is from a clay-deficient area where boron has not already been removed in significant amounts. This situation requires flow into a shale section, probably from a sandy, more porous medium, and is not a very likely hydrodynamic situation, especially for the amount of flow needed. If boron is not primarily gained from an external source (diffusion also does not appear to

286

K.H. WOLF AND G.V. CHILINGARIAN

be reasonable mechanism on any large scale), one must look for a redistribution within the shale so that boron can be added to the lattice of the illitic material forming diagenetically. A newly deposited sediment will have boron located in two regimes: boron adsorbed on the clay surfaces at the depositional site and previously inherited ‘detrital’ boron in the clay lattice. The adsorbed boron would then be retained through burial and finally incorporated into the lattice of the new illitic material formed during diagenesis.” Perry (p. 159) finally concluded that the boron-paleosalinity technique may only be useful as a very general indicator of depositional environment, and only after careful clay mineral determination and petrologic interpretations. Couch also stated in his personal communication that certain corrections explained by him have to be applied in order to give paleosalinity values which are in good agreement with other evidence. He concluded that as long as one is concerned with rocks of low permeability, postdepositional processes apparently do not alter the original boronsalinity relation. Walker (1964)discussed the use of the boron content of sediments as a paleosalinity criterion and concluded that, in spite of numerous complexities which demand a careful scrutiny and adjustment of the boron data obtained, this geochemical tool does supply a reliable measurement of paleosalinity. Some of these complexities described by Walker are related to the postdepositional enrichment of boron. He remarked (p. 211 and pp. 215-216) that the coarse-grained sandstones are commonly porous and have an anomalously high boron content. If compaction fluids from shales in more saline parts of the basin moved into the porous sandstone bodies, a reaction between the clay minerals (especially illite) in the sandstones and the boron might occur. The boron content of normal sea water (salinity of 30-40 parts per thousand) is approximately 50 ppm, whereas the observed boron content of marine illites is 200-300 ppm. This requires an equilibrium in an open system between the sea water and the illite and, of course, a slow rate of sediment accumulation. On the other hand, in a closed system where the clays are exposed to trapped interstitial fluids with a low boron content of around 5 ppm, a small enrichment of boron in the illite within the porous sandstone would result in a large relative decrease in the boron concentration of the interstitial solution. Consequently, there would be no appreciable enrichment of the boron in the illite. All of the above-considered publications refer to the important influence subsurface fluids may have on boron content. Some authors even mentioned compaction fluids as a possible variable directly controlling the boron concentration. In general, whenever various researchers refer to interstitial waters or subsurface fluids without being more specific, it is safe to assume that at least part of these solutions may be compaction waters. Whenever the data allows, one should be more specific as to the precise origin of the fluids.

DIAGENESIS OF SANDSTONES AND COMPACTION

287

Other trace-element contents Although the above considerations are confined to boron in sediments, similar allowances must be made t o the possible presence or absence of the effects of compaction fluids on the content of any other trace element, e.g., potassium, sodium, uranium, copper, lead, and zinc, to name only a few. Ericsson (1972),in his publication on the chlorinity of clays as a criterion for paleosalinity, remarked that the measured chlorine contents often varied greatly. He stated (p. 5) that grain size and organic carbon content have some influence on the vertical chlorinity changes within sediments. The vertical salinity variation in a particular area may also be influenced by diffusion and leaching, either by surface or subsurface water. These processes are controlled, for example, by compaction, but the magnitudes of various effects are not known. In addition, salinity of compaction fluids changes upon migration (Rieke and Chilingarian, 1974,pp. 239-272). Hartmann (1963)described, among other interdependencies of chemical elements, the relationship between iron and vanadium in sandstones. It is reviewed here as an example of an element-element correlation. Both Fe and V behave similarly geochemically, which results in their parallel occurrence in (a) hematite-rich, clay-poor sandstones (Fig. 3-175),and (b) clayrich, hematite-free sandstones (Fig. 3-176).The former case is more or less self-explanatory, because the vanadium is associated with the iron in hematite as a consequence of their similar geochemical behavior. In case b, the iron associated with the clay minerals was at the same time accompanied by vanadium, as has been demonstrated in the laboratory. On assuming that fluids, compaction waters or otherwise, can precipitate authigenically new

Fe,O,

CONTENT, %

Fig. 3-175. Graph demonstrating the dependency of the vanadium content (ppm V) on the total F e z 0 3 content (%) in red, hematitic, clay-poor sandstones. (After Hartmann, 1963, fig. 6; courtesy Geochim. Cosmochim. Acta.)

K.H. WOLF AND G.V. CHILINGARIAN

288

h,O,

CONTENT, Vn

Fig. 3-176. Graph demonstrating the dependency of the vanadium content (ppm V) in the total F e z 0 3 content ( W )in hematite-free sandstones and claystones. (After Hartmann, 1963, fig. 5; courtesy Geochim. Cosmochim. Acta.)

minerals (such as hematite), can cause alterations (e.g., hematite to goethite, etc.), and can dissolve minerals, then the occurrence of any accompanying minor and trace elements will exhibit certain regularities, whether they occur adsorbed on the surfaces of minerals or form part of their internal crystalline structure. Compaction fluids and isotope composition

In regard to the possibility that compaction fluids may be responsible for determining the final isotope composition of sediments, one example is considered here. Monster (1972,’ p. 941) stated that “. . . there is no apparent reason to assume that in a particular crude oil the sulfur in the sulfur-bearing organic compounds has a uniform isotopic composition. On the contrary, from what is known of the sulfur isotopic composition of organically bound sulfur in marine sediments, there are good reasons to expect considerable heterogeneity in sulfur isotope ratios.” Studies of recent sediments have revealed a large isotopic heterogeneity for sulfur, including organically bound sulfur, with depth and location within a particular basin. On the other hand, there was a certain S-isotope homogeneity for various oil fractions obtained by Monster. This is in strong contrast to the wide variability in isotopic content of the organic sulfur compounds in recent marine sediments, which are supposedly the precursors of petroleum in reservoir rocks. There are two processes, in particular, that can effectively increase the S-isotopic homogeneity: (1) primary migrational mixing, and (2) S-isotope exchange reactions.

DIAGENESIS OF SANDSTONES AND COMPACTION

289

(1)It is reasonable to assume that any particular basin has a wide range of isotope ratios for organically bound sulfur, whereas the range in ratios will be relatively small at any particular location. The mixing of the various organic sulfur compounds in the fluids that undergo either lateral or vertical migration, commonly caused by compaction, may result in a reduction of the range of isotope ratios. At the final location, each individual chemical compound will have an averaged isotope composition resulting from the contributions of various locations. (2) Exchange reactions between different compounds can only take place if they are in proximity to each other. Again, compaction fluids can achieve transportation of various compounds to bring them together for chemical reactions. In this regard, Monster (1972,p. 946) stated that many chemical reactions will take place during diagenesis, migration, and maturation. According to him, "addition of extraneous sulfur compounds at some stage of migration will, in general, change the 34S/32S ratio of total sulfur of the oil, and might extend the range of isotopic ratios". Similar arguments can be applied to C-isotopes. Figures 3-177 and 3-178 are schematic diagrams indicating the major processes affecting the isotopic composition of petroleum moving from a primary source to a reservoir rock. The values presented apply only to a particular marine sedimentary basin; -40

GAS PHASE -30 to-38

-?lo

&

6c'3%0

- 20

-I

I0

,

0 I

I

,

I

,d

LlOUlD PHASE - 2 0 to -23

Fig. 3-177. 3C/ 2C in petroleum from primary source to reservoir rock in typical marine sedimentary basin. (After Monster, 1972, fig. 3, p. 947; courtesy Am. Assoc. Pet. Geologists.)

290

K.H. WOLF AND G.V. CHILINGARIAN

they are not general, world-wide ranges, so that similar diagrams for other marine basins may vary in range. Additional complexities can be visualized if one considers isotope variations in the (a) normal, (b) evaporitic, ( c ) barred (i.e., "Black-Sea" type), (d) brackish-water, and (e) fresh-water basins. Certain geologic conditions could result in mixing of two or more of the fluids derived from these environments during movements of compaction fluids in the subsurface. FLUID-DIAGENESIS IN SANDSTONES

The field of fluid-diagenesis is a comparatively new one in contrast to the study of diagenetic textures and minerals. The latter may reflect chemical changes in intrastratal fluids, which caused the precipitation of cements and led t o other secondary alterations. It is easier to investigate the visual effects of diagenetic processes and use them for theoretical interpretation of the causes. With so many new precise instruments and techniques now at hand, however, there is a trend in investigating the geochemical and physical con-

DIAGENESIS OF SANDSTONES AND COMPACTION

291

ditions that resulted in diagenetic modifications. This can be done by studying recent sedimentary environments and by laboratory experiments and then using the results for establishing “diagenetic conceptual models” for a spectrum of environments in nature. For example, precise investigations of fluids under natural conditions as well as in vitro promise to provide data on the interactions of fluids and sedimentary particles. Inasmuch as compaction fluids are believed to constitute one major group of solutions responsible for diagenesis, a brief discussion is presented here. In general, fluid-diagenesis in sediments has been conveniently divided according to the solid component involved in the chemical reactions, e.g., fluid-diagenesis in clayey, sandy, and pyroclastic deposits. As has been pointed out in other sections, as a result of the hybrid origin and composition of sandy sediments, the data available on fluid-diagenesis in clayey sediments, for example, is applicable also to clayey sandstones. Runnels (1969)discussed an interesting aspect of fluid-diagenesis, namely, the precipitation of minerals as a result of mixing of two different solutions. Numerous varieties of natural waters are summarized in Table 3-XLVIII, and inasmuch as sediments accumulate in each one of these environments, the trapped primary interstitial fluids vary in composition accordingly. The secondary differential fluid alterations during burial complicate the situation further. Many of these sedimentary facies are closely associated in three dimensions, so that during migration as a result of compaction the various types of fluids will mix (see also Shelton, 1964). Runnels (p. 1188)mentioned some of the consequences: (a) precipitation of minerals; (b) production of porosity through mixing of sodium chloride waters (he plotted the solubility of calcite as a function of the content of neutral electrolyte, NaC1, TABLE 3-XLVIII Classification of natural waters (after Runnels, 1969, table 1 , p; 1189) I. Meteoric waters

11. Surface waters A. Lakes: (1) fresh, ( 2 ) saline B. Streams: ( 1 ) fresh, ( 2 ) saline C. Swamps: (1) fresh, ( 2 ) saline D. Oceans 111. Subsurface waters A. Vadose: ( 1 ) fresh, ( 2 ) saline B. Phreatic (1) normal formation water: (a) fresh, (b) saline ( 2 ) ascending thermal water: (a) fresh, (b) saline

K.H. WOLF AND G.V. CHILINGARIAN

292

in solution); and (c) dissolution (e.g., decementation) and precipitation (e.g., cementation). He showed that an understanding of these processes is very important in petroleum-reservoir and ore-genesis studies (e.g., Colorado Plateau uranium ores; fluorspar of the Kentucky-Illinois district; and Mississippi-Valley-type Pb-Zn ores). The simple classification of natural waters, as given in Table 3-XLVIII, incorporates 14 separate categories of waters, which in turn can give 91 combinations on mixing any pair of fluids. When one considers, however, that in nature three, four or even more fluids can mix in both surface and subsurface environments, then many more combinations can be envisaged. The complexity of fluid geochemistry is even greater when other parameters are considered, e.g., temperature and pressure differences and chemical variations with depth, which have been established for many sedimentary basins. Runnels considered the upper limit of diagenesis to roughly correspond to conditions of temperatures and lithostatic pressures in the deepest wells drilled (20,000-25,000 f t depth), i.e., temperature of about 200°C, pressure of 1500 atm, temperature gradient of loC/10O ft, and pressure gradient of 100 psi/lOO f t for water-saturated sedimentary rocks. Under the more extreme conditions, the dissociation of water as a function of temperature and pressure (Blatt et al., 1972) must be considered in geochemical investigations on fluiddiagenesis (Fig. 3-179). Sharma (1968, p. 232) stated that petrographic investigations are the study of the end results or effects of chemical reactions involving pore fluids and the solid matter of the host material, so that petrographic data can give

5e 0

80

40

I20

TEMPERATURE, I

!moo

I

~qaoo

200

160

1

1

240

OC 1

1

2 o w 3opoo

DEPTH, ft

Fig. 3-179. Variation in the dissociation constant of water as a function of temperature and depth in a geosyncline. (After Blatt et al., 1972, fig. 7-1; copyright 0 1972 PrenticeHall, Englewood Cliffs, N.J.)

DIAGENESIS OF SANDSTONES AND COMPACTION

293 THIN PLASTIC

Fig. 3-180. Sharma’s sediment model. (After Sharma, 1968, fig. 1, p. 233; Minemliu m Deposita. )

rtesy

only a partial answer to the questions on the origin of diagenetic products. To fill this gap, he performed experiments to determine: (1)the relationship between the chemistry of the water in the sediments and the mineralogy of the sedimentary particles; (2)the factors controlling the processes of calcite and silica precipitation in sediments; (3) the processes responsible for transforming sediments into hard rocks; and (4)the paragenesis involved. Sharma used a box containing various sediments, as shown in Fig. 3-180, and pumped fluids of known composition (i.e., artificial sea water of different compositions) through the sediments. Chemical analyses were performed on the effluent waters collected at regular intervals, whereas petrographic and X-ray analyses were done on the chemical precipitates formed during the experiments. Table 3-XLIX shows the chemical changes that took place. Both aragonite and silica were precipitated as cements of various forms. Sharma utilized a total volume of solutions of approx. 4500 ml in his model experiments with a consequent flow of 35.5 ml/cm2. This requires a burial depth of 47 cm, if the data by Emery and Rittenberg (1952,p. 755)is accepted. These authors stated that a burial of 100 cm of sediments will result in a flow of interstitial fluids of 75 ml/cm2 due t o compaction. Thus, the experiments performed were comparable to situations of early diagenesis during shallow burial. Sharma (1970) discussed: (1)diagenetic alterations of montmorillonite and degraded illite to chlorite, possibly by ionic influence in the marine

K.H. WOLF AND G.V. CHILINGARIAN

294

TABLE 3-XLIX Chemical composition of influent1 and effluent water injected through the model in run No. 2 (after Sharma, 1968, table 2, p. 235) 7/25/63 Influent water ~

Potassium Magnesium Calcium Strontium Lithium Chloride Sulfate Bicarbonate Carbonate Specific gravity Resistivity PH

310 ppm 651 ppm 128 ppm 11 PPm 4 PPm 21,780 ppm 1240 ppm 24 PPm 48 PPm 1.028 0.312 ohm/m 9.01

7/29/63 Effluent water

Gain

Loss

~

290 ppm 749 ppm 528 ppm 32 PPm 3 ppm 21,080 ppm 1359 ppm 1579 ppm 0 PPm

98 PPm 400 ppm 21 PPm 119 ppm 1555 ppm

1.027 0.312 ohm/m 6.95

20 PPm

1 PPm 700 ppm

48 PPm

0.001

The influent water was injected with a 90-psi carbon dioxide pressure at 72"F.

environment, and (2) changes of kaolinite t o illite or chlorite, which are often present in sandstone matrices. These mineral changes distinctly alter the chemistry of the trapped pore sea water, which often exhibits vertical variations in composition in sedimentary sections. As pointed out by Sharma (1970, p. 722), under ideal conditions the intrastratal fluid composition may be controlled by the physical characteristics (grain size, packing, porosity, permeability, etc.) and the composition of the mineral phases with which the fluids are in contact. The degree and rate of compaction are also very important. Buried waters are usually different, in composition, from the overlying sea water as has been demonstrated by numerous investigators, e.g., by Friedman et al. (1968) and Friedman and Gavish (1970). Hydrodiagenesis depends on: (1) changes in the depositional environment; (2) rate of sediment accumulation; (3) length of exposure of detritus to fluids; (4) mineralogy of the sediments; ( 5 ) differential ionic movements due to diffusion and semi-permeable effects of clays; and (6) cation exchange. All of these factors are related to compaction and compaction fluids. For example, slow rate of deposition prolongs watersediment interaction and reduces the rate of compaction. There is also a longer interaction at the sediment-water interface as well as a prolonged subsurface watersediment interaction under these conditions.

DIAGENESIS OF SANDSTONES AND COMPACTION

295

Sharma’s (1970) investigation concentrated on the glacial sediments that were transported from Alaska t o the sea. The glacial material was mechanically eroded and abraded, but underwent a minimum amount of chemical alteration (i.e., chemical weathering was absent) on land prior t o being deposited in the sea. Inasmuch as these glacial sediments are chemically the most immature, the chemical interaction between the sedimentary particles and the sea water and pore fluids would be expected t o be particularly distinct. For example, hornblende and biotite release K+ and Mg2+ ions into the fluids. Sharma (1970, p. 727) stated that the initial composition of intrastratal fluids is controlled by that of the depositional milieu, such as type of surface water, presence and degree of flocculation of clays and organic matter, and ion exchange. Alterations in both surface and shallow interstitial fluids can be seasonal in fjords, estuaries, saline lagoons, and lakes, among others. The mineralogy of sediments, proportions of sand, silt, and clay fractions, and the chemistry of the interstitial fluids would determine the types of chemical reactions. For that reason, it was so important to use very immature mineral particles in Sharma’s investigation to eliminate at least one variable, namely, chemical alterations of the solid particles as a result of surface weathering prior to erosion and transportation. Grain size and degree of compaction is significant in controlling diffusion. According to Sharma (p. 729), the diffusion coefficient of dissolved ions or molecules in water is of the order of l o v 5 cm2/sec., but when the free ions or molecules encounter sediments, their movement through the pore fluids is impeded by the sedimentary particles. (See also Peck, 1967, for discussion on diffusion coefficients.) Duursma (1966a,b), among others, showed that many heavy metals, alkaline earths, and even some alkali metals, have a great affinity for clayey calcareous sediments, so that the ionic diffusion coefficient is reduced by factors up to lo6. This indicates that diffusion over a few centimeters may take several years, but if compaction is considered, the rate of fluid movement is increased. In sediments that contain clay minerals, some factors and processes have to be considered that would not be important in pure sandstones. Clays possess the properties of semi-permeable membranes resulting in salt-sieving or reverse osmosis, e.g., when a saline fluid is passed through a layer of clay minerals, some salts remain behind. This complex process is determined by fluid chemistry, porosity and permeability, degree and rate of compaction, burial depth (i.e., overburden pressure), composition of sediments, and possibly temperature (i.e., geothermal gradient). Where studies are undertaken on near-surface sediments, as done by Sharma, the absence of significant overburden and compaction, therefore, will reduce the amount of interstitial

K.H. WOLF AND G.V. CHILINGARIAN

296

fluid movement and the membrane effects will be minimized. Slow compaction, as a result of slow sediment accumulation, would indicate that sedimentary particles were exposed longer to surface fluids and that, at the same time, compaction fluids evolved slower. In the absence of compaction, the sedimentary particles are not exposed to compaction fluids moving from below. Dickey (1972)pointed out that chemical processes associated with compaction result in mineralogical and other chemical changes of importance in both petroleum and ore studies. With increasing depths and geologic time, the pore fluids increasingly deviate from the composition of sea water: sulfate and bicarbonate anions are usually lost, chloride anion is left behind, Ca2+ cation content is increased and Mg2+ is reduced (3 to 5 times as much Ca2+ as Mg2+, whereas in sea water the Mg/Ca ratio is 3/1). Variations in fluid composition with depth, however, are very common, as exemplified by Fig. 3-181 (Schmidt, 1971). The pore fluids of shales or 0

1

I

E - I I-

n w 3000

-

XXJOI 0

I

50

I

IM)

Is0

1

200

TOTAL DISSOLVED SOLIDS, g/l

Fig. 3-181. Variations in concentration of interstitial water with depth, southwest Louisiana. The water in the pores of the shale is much less concentrated than that in the adjacent sands. Water in the sands of the high-pressure zone at Manchester is much less concentrated than that in the equivalent normal-pressure sands at Hackberry. A = water produced from sand, Hackberry field (normal pressure); B = water extracted from shale, Manchester field; C = water produced from sand, Manchester field (abnormally high pressure); D = low-density and high-pressure shale. (After Schmidt, 1971; in: Dickey, 1972,fig. 11, p. 11;courtesy Int. Geol. Congr., Montreal.)

DIAGENESIS OF SANDSTONES AND COMPACTION

so;

29 7

HCO;

Fig. 3-182. Triangular diagram showing the relative amounts of anions in interstitial waters in formations of southwest Louisiana. The waters in the sands contain almost no bicarbonate and sulfate, but these ions predominate in the shale waters. A = Hackberry sand waters (normally pressured); B = Manchester sand waters (abnormally pressured);C = normally-pressured shale water; D = abnormally-pressuredshale water. (After Schmidt, 1971, in Dickey, 1972, fig. 12, p. 12; courtesy Int. Geol. Congr., Montreal.)

muddy sediments are quite different in composition from the solutions in the associated permeable sands, as demonstrated in Fig. 3-182for one particular area: the sandstones contain mainly chloride and are poor in or devoid of sulfate and bicarbonate anions, whereas clayey sediments contain the latter two anions predominantly. The concentration of the solutions in sandstones are related to the degree of compaction of the adjacent clayey deposits, and the water samples taken from the high-pressure zones are often less concentrated in dissolved solids than those from normally pressured zones. Abnormal subsurface pressures are found only in sandstones completely enclosed in shale, with no permeable connection t o the outcrop or into adjacent porous units. According to Dickey (1972),the fluids from high-pressured zones often have salt concentrations much less than is normal for their depth of burial. Dickey suggested that the processes which caused the concentration of the interstitial solutions, were arrested at about the same stage at which the compaction of the clayey sediments was terminated. Dickey stated that on a regional scale, there is a general tendency for the concentration of subsurface fluids to increase with depth, with possible more local deviations and reversals. Von Engelhardt (1960)also commented on the general trend: drilling into sedimentary basins has shown that the salinity of pore fluids commonly increases with depth. This is easy to observe on electrical resistivity curves as exemplified in Fig. 3-183.Another example is given in Fig. 3-184,where at a

298

K.H. WOLF AND G.V. CHILINGARIAN

SP -4

-+

DEmm

RESISTIVITY

W mV

0

C

D

-I

m

i

D

2

D II

<

+

C D 0

m D -4

m II

2

t

2

C

0

D

rn D

-I

rn D

-I P D

< -

Fig. 3-183.Resistivity and SP measurements using the Schlumberger method in a borehole near Stockstadt (Upper Rhine region), Germany. The decrease of the resistivity between 300 and 400 m and the corresponding change in the SP indicate the increase of the salt content of the pore solutions. (After Von Engelhardt, 1960,fig. 59;courtesy Springer, Berlin.) Fig. 3-184.Salt content (96)as related to depth of burial of the formation fluids of the St. Genevieve Sandstone (Mississippian) in the Illinois basin (according to 235 analyses by Meents and coworkers). (After Von Engelhardt, 1960,fig. 60; courtesy Springer, Berlin.)

depth interval of 300-400 m the salt content increases rapidly with depth (Meents et al., 1956). Rieke and Chilingarian (1974),however, found so many exceptions that they are opposed to such a generalization.

DIAGENESIS OF SANDSTONES AND COMPACTION

299

Schmidt (1973)presented data on interstitial fluid composition and geochemistry in a shale-sandstone section with abnormally high pressure zones, and noted significant differences between the total dissolved solids concentrated in waters from normally pressured sandstones (600-180,000 mg/l) and those from overpressured sandstones (16,000-26,000 mg/l). He stated (p. 321): “. . . conversion from predominantly expandable t o non-expandable clays accelerates near the top of the high-pressure zone, which appears correlative with a major temperature gradient change, an increase in shale porosity (decrease in shale density), a lithology change to a massive shale, an increase in shale conductivity, an increase in fluid pressure, and a decrease in the salinity of the interstitial waters.” In this case, decrease in water salinity can be due t o release of water during conversion of montmorillonite to illite and/or compaction effects, which are described in detail by Rieke and Chilingarian (1974). Zimmerle (1963)differentiated between “destructive” and “constructive” diagenesis. In the former case, the complex minerals change to less complex ones or are replaced by other minerals, and this type of diagenesis does not lead to reduction of porosity. Constructive diagenesis refers to overgrowth or authigenic neomorphism of new, commonly simple minerals, usually in connection with chemical mobilization or addition of material. This results in a decrease in porosity and permeability, often related t o mechanical compaction, movements of compaction fluids, and pressure solution. The latter may deliver the chemical components, so that compaction, pressure solution, and cementation can be in some cases interdependent. After describing the various diagenetic features, Zimmerle offered a paragenetic interpretation based on pH of the solutions (Fig. 3-185).Solutions are the “carriers” of the chemical compounds and the dissolution and precipitation of minerals are dependent on many variables. Based on the work by Packham and Crook (1960)and others, however, Zimmerle used pH as the main parameter in Fig. 3-185. Siderite, chamosite and glauconite are of synsedimentary origin under neutral pH and medium values of Eh. Pyrite formed under reducing and neutral t o weak acidic conditions during early diagenesis. It formed prior to neomorphic quartz and feldspar, as indicated in thin sections. Kaolinite, replacing primary clays and clastic feldspar, and neomorphic brookite also formed prior to quartz precipitation under acid pH values. Secondary quartz, which is younger than brookite and older than calcite, indicates a weak acidic to neutral pH, the feldspars (i.e., albite, orthoclase, or microcline) suggest a neutral to weak basic milieu, and calcite requires a basic pH value for its precipitation. A second generation of quartz and kaolinite, apparently related to the introduction of oil into the rock, was followed by barite: It should be pointed out here that movements of petroleum are often considered to be the result of compaction of fine-grained

300

K.H. WOLF AND G.V. CHILINGARIAN PH RANGE.

SEOUENCE

F g $ #--

Syndloqmettc Early Diapcnctic

Late

Diopmcfic

I.

A

2.

m

Fig. 3-185. Schematic paragenesis of the cement and the neomorphic minerals depending on the maximum pH value in the Dogger-fl Sandstone of the Plan-Ost Oil Field. 1 = pH of the formation waters = 5.9-6.6; 2 = range of the “intrastratal solutions” for which evidence is available. (After Zimmerle, 1963, fig. 8, p. 15; courtesy Erdol Kohle.)

sediments. From the paragenesis, Zimmerle deduced that the fluids in the sandstone had a pH range of 5.9-6.6, i.e., weakly acidic, and concluded that both depth of burial and fluid composition were important in determining the sequence of mineral precipitation. During diagenesis a “reversal’,’of porosity occurs, i.e., the originally clay-poor and porous sandstones become calcite- and silica-cemented with elimination of pores, whereas the originally clay-rich, less porous sandstones are influenced only slightly by diagenetic processes and maintain most of the original minor porosity. Hence, in studies of the history of porosities, not only compaction history but also the development of cementation is of significance. Sharma (1969) presented a pH-paragenetic diagram indicating changes with depth of burial (Fig. 3-186),similar to the one offered by Zimmerle (1963). Potter (1968) showed an influence of intrastratal solution on garnet and concluded that for”ancient indurated or strongly compacted sediments the use of heavy minerals for stratigraphic-correlation and provenance studies

301

DIAGENESIS OF SANDSTONES AND COMPACTION

Fig. 3-186.Diagram showing pH relations during all stages of paragenesis in Halfway Formation as observed in cores. Order of dissolution and precipitation of minerals is also presented. (After Sharma, 1969,fig. 9;courtesy Minemliurn Deposita.)

may be of limited value, because these sediments have been corroded and removed. Inasmuch as the survival of minerals depends on a number of interrelated factors, e.g., pH of solution, nature of minerals involved, permeability and porosity of beds, temperature, pressure, rate of flow of fluids, and total geologic time, it is evident that the degree and rate of compaction and the characteristics of the compaction fluids would be influential in determining the corrosion of heavy minerals. Inasmuch as bacteria are extremely important in diagenetic alterations of sediments, it should also be pointed out here that compaction, which modifies the characteristics of the sediments, also alters the number of bacteria and, possibly, the type of bacterial assemblage, bacterial migration, and bioTABLE 3-L Distribution of bacteria in sediments according to particle size (after ZoBell, 1943,in Strakhov, vol. 2, 1969,table 41,p. 424) Sediment

Average diameter of sedimentary particle (p)

Content of N,

Water content

(%I

Number of bacteria per g of dry residue (in thousands)

Sand Silt Clay Colloids

50-1000 5-50 1-5 <1

0.09 0.19 0.37 1.0

33 56 82 98

22 78 390 1500

(%)

MOISTURE CONTENT,'/t

Fig. 3-187. Change in moisture content in sediment cores from the Bering Sea. (After

A.P. Lisitzyn, in Strakhov, 1969, fig. 181, p. 425.)

DIAGENESIS OF SANDSTONES AND COMPACTION

0

I

LOGARITHM OF NUMBER OF BACTERIA 3

5

7

303

9

E u

z I f 60 k

80

Fig. 3-188. Vertical distribution of bacteria (number of bacteria per g of wet mud) in oceanic sediments. (After ZoBell, 1946, in Strakhov, 1969, fig. 182, p. 425.)

genic processes. The factors affecting bacterial growth include: (1)grain size of the sediments (Table 3-L); (2) moisture content of the sediments (Fig. 3-187); (3) content of organic material (Table 3-L); (4) depth of burial (Fig. 3-188); (5) degree of compaction. Inasmuch as the above parameters are interrelated, compaction influences diagenesis directly by modifying the bacterial activity. The literature on the interaction of interstitial fluids and volcanic detritus is becoming increasingly voluminous, in particular on the processes related to zeolite genesis (e.g., Hay, 1966; Utada, 1968; Sheppard and Gude, 1968; Mariner and Surdam, 1970; Utada and Minato, 1971; Ijima and Utada, 1971; Sheppard, 1971; Surdam and Parker, 1972). Specific references to compaction fluids within pyroclastic units and their influence on diagenesis and burial metamorphism, however, are less frequent, probably because the results of compaction-fluid reactions are difficult to nearly impossible to distinguish at the present time from the effects of reactions of any other type of intrastratal solutions. Nevertheless, occasionally early researchers have considered compaction in pyroclastics, and recent work on the compaction of volcanic debris has been treated in Chapter 6 of this book. As early as 1954, Coombs observed in the Triassic succession of New Zealand fairly normal sedimentary diagenetic phenomena, such as overgrowths on feldspars, cementation by quartz and chloritic minerals, and zeolite replacement of volcanic glass. He stated that in addition to connate waters trapped in the sedimentary pile, large quantities of fluid were stored

K.H. WOLF AND G.V. CHILINGARIAN

304

up in the volcanic glass and in the zeolites of early origin. As the temperature during increasing burial rose, perhaps up to 150-300°C at the base of the pile, this stored-up water facilitated a special type of metamorphism with some low-grade hydrothermal effects. Dickinson (1962a,b) discussed diagenetic to epigenetic alterations in mineralogy and chemistry of pyroclastic material and concluded that the metasomatic processes causing zeolitization, phyllosilicatization, and albitization were not initiated by fluids from an igneous source. Instead, the nearest

IN I T l A L

FINAL Water flows through membrane

(A1

semi-permeable membmne transmits water but not salt m re less equal concentrations Concentrated

equal concentrations

concentrated

-

Water flows upward through shale membrane

(C

by hydrostatic head

Fig. 3-189. Initial and final equilibrium states 01 normal (A) and reverse osmosis (B) in a U-tube. C. Reverse osmosis in two sandstones separated by a thin shale acting as a semi-permeable membrane. (After Pettijohn et al., 1972, fig. 10-9, p. 413; courtesy Springer, New York.)

DIAGENESIS OF SANDSTONES AND COMPACTION

01

00

I

80

I

1

1

60

40

20

305

OIL SATURATION,%

Fig. 3-190. Capillary pressure curves (displacement of water by oil): I = coarse-grained sandstone; II = clayey fine-grained sandstone. (After Von Engelhardt, 1960, fig. 34; courtesy Springer, Berlin.)

source of reactive fluids is formed by the connate pore waters and adsorbed aqueous films that must have been expressed from associated and underlying marine mudstones by compaction. Hay and Ijiama (1968) briefly considered compaction by remarking that the tuffs of Hawaii show no compaction features and that the net increase in bulk density observed by them from younger to older tuffs, seems to be a reflection of a net increase in constituents precipitated by inter-pore fluids. Pettijohn et al. (1972, p. 414) discussed a special process, i.e., reverse osmosis, that must also be operative during compaction of sediments. This process that affects the composition of subsurface waters by “salt filtering through a semi-permeable membrane” depends on the content of clayey sediments (see De Sitter, 1947, Bredehoeft et al., 1963, and Rieke and Chilingarian, 1974). The differential pressure at depth within the sedimentary pile supplies the energy. Figure 3-189,a illustrates that in simple osmosis the difference in concentration of salt in fluids on the two sides of the membrane results in a pressure differential (= osmotic pressure), which causes the movement of the solvent from the low-concentration to the highconcentration side. In “reverse osmosis”, the same concentration of salt exists on both sides of the membrane, but a pressure differential is formed that induces a flow of the solvent from the high- to the low-pressure side and results in a higher salt concentration on the high-pressure side (Fig. 3-189,b). According to Pettijohn et al. (1972), applying these principles to a sedimentary environment, the pressure differential with varying depth of overburden

K.H. WOLF AND G.V. CHILINGARIAN

306

will lead to a situation as shown in Fig. 3-189,c. Here, a sandstone below a shale acting as a semi-permeable membrane will have a higher concentration of salts than a sandstone above that shale. It is difficult t o prove at the present stage of knowledge, however, to what degree “salt filtration through a semi-permeable membrane” process is actually occurring in sedimentary basins. A number of important fields of study related t o diagenesis and fluid movements should be given more attention in the future by sedimentary petrologists. These include: (1) capillary pressures and their relationships to pore geometry (see Chilingarian et al., 1972); (2) relative permeabilities of sediments to oil, gas, and solutions of different compositions at varying temperatures and pressures (see Langnes et al., 1972); and (3) relative degrees of adsorption of fluids of different compositions on grain surfaces. Figure 3-190 shows capillary pressure curves for a coarse-grained sandstone and for a clayey, fine-grained sandstone, applied to conditions required for water t o replace oil. Figure 3-191 indicates the relative permeabilities to oil and to water, where water is the wetting agent and oil is the non-wetting phase. The effects of interface phenomena on diagenesis will have to be considered thoroughly in future research, because these phenomena affect the microscopic and sub-microscopic processes during diagenesis in coarsegrained sediments. Several examples have been discussed in this chapter which demonstrate that the presence of a film of oil on sand grains suppress

0

20

40

80

60

WATER SATURATION,

100

o ?‘

Fig. 3-191. Relative permeabilities of a rock to oil (kro) and to water (kr,,,). Water is the wetting phase, whereas oil is the non-adsorbed phase. (After Von Engelhardt, 1960, fig. 36; courtesy Springer. Berlin.)

DIAGENESIS OF SANDSTONES AND COMPACTION

307

-

#

20-

c

UJ

d w

s I-

n W

m

IY

0

m

n l0-

y.'

' .

0

.

INTERNAL SURFACE AREA.S/E..IO'

. 0.mcm a

Fig. 3-192. Amount of adsorbed water (Sh, on surfaces of the pores) and internal surface of 7 5 samples of the Bentheimer Sandstone (Valendis; Scheerhorn Oil Field), area (S/€) Germany. (After Von Engelhardt, 1960, fig. 37; courtesy Springer, Berlin.) Fig. 3-193.a. Curvature of the wetting phase (ring-like) at the point of contact between two spheres (schematic). (After Von Engelhardt, 1960, fig. 38; courtesy Springer, Berlin.) b. Calculated profiles of the drop at the contact of two spheres considering different relationships between the capillary pressure ( P c ) and interfacial tension (0).(After Von Engelhardt, 1960, fig. 39; courtesy Springer, Berlin.)

the chemical interaction of the grains with solutions. The type of distribution of adsorbed water in pore spaces has been rarely observed directly, so that most conclusions are based on indirect considerations. It has been theoretically determined that there is a linear relationship between the surface area and the amount of adsorbed water, which can be substantiated by measurements on natural rocks, e.g., see Fig. 3-192. Such

K.H. WOLF AND G.V. CHILINGARIAN

308

data is of value in diffusion studies of chemical elements during diagenesis. Figure 3-193shows the relationship between capillary pressure ( P c ) ,surface tension ( o ) , and the curvature of adsorbed water rings at the contacts between two grains. More precise work on the relationships of capillary water, its movement patterns, pressures required to move it, shapes of the water films, diffusion, and related factors t o precipitation of minerals in the interstitial spaces is needed. It has already been suggested, for instance, that cement occurring in the textural relationship to its host grains as does the adsorbed water in Fig. 3-193,is the result of precipitation from such adsorbed solutions. The amount of fluid available from one such adsorbed ring, however, would be insufficient to account for the same amount of cement; therefore, it must be assumed that additional volumes of chemical elements were supplied by diffusion or migration of solutions. Magara (1974) concluded that fluid movements from shales into associated sands are the result of both compaction and osmosis, as schematically illustrated in Fig. 3-194.During compaction, the water from the clayey sediments (now shales) will be squeezed into the more permeable sands, as the clay-rich deposits will decrease in porosity accordingly. Maximum reduction in porosity or fluid expulsion will occur in the clayey sediments directly above and below the sands, whereas the porosity near the center of the shale will remain relatively high, but subsequently will be affected (reduced) with increasing compaction. The fluid pressure as a result of compaction can be A I ALE POROSJM

8

FLUID PRESSURE IN WALE

C WATER SALINITY IN SHALE

SHALE

SHALE

SHALE

-.#gfifiw.s* ............. SHALE

owm

.:::::.z.k.:.:.:.: +d .....:L :: Di"" ...............

1 ................

.:.:.:.>>w:p$:::::: ::::::::*.A+:.:.:.:.:.

T

LHYDROSTATIC PRESSURE Fig. 3-194. Schematic diagram showing shale-porosity, fluid pressure, and pore-water salinity distributions in interbedded sand-hale sequence. (After Magara, 1974, fig. 6, p. 288; courtesy Am. Assoc. Pet. Geologists,)

DIAGENESIS OF SANDSTONES AND COMPACTION

309

calculated from the shale-porosity distribution, as done by Magara (1968, 1969). Figure 3-194,B shows schematically the fluid pressure plotted corresponding to the porosity distribution. As expected, the water in the clayey deposit will move from a zone of higher, excessive pressure to a lower-pressure zone. Compaction-water movements are indicated by arrows in Fig. 3-194,B. As a result of ion-filtration, the ions are concentrated in the clayey unit, as schematically illustrated in Fig. 3-194,C. Salinity is the reciprocal of porosity, in that it increases as the porosity decreases. Thus, the salinity in the shales increases toward the sands. The process of osmosis induces water to move from a fresher to a more concentrated locality, as indicated in Fig. 3-194,C, but the osmotic pressure difference in this case is not very pronounced in contrast to that due to compaction. The flow of solution due to the combination of compaction and osmosis will continue until the clay-rich units reach equilibrium, and no fluids can be expelled from them by compaction. Salinity also may reach equilibrium. If, on the other hand, a “freshening mechanism”, such as dehydration of montmorillonite, changes the salinity later on, the osmotic fluid may be changed. Magara (1969, p. 289) stated, however, that the most important in this combined mechanism of flow is that the salinity contrast resulting from ion-filtration starts to appear during early compaction. Consequently, the resultant osmotic pressure difference seems to support fluid migration from the clayey sediments at the early stages of water expulsion. In their thorough discussions on the effect of compaction on salinity, Rieke and Chilingarian (1974) reached different conclusions, e.g.: (1)salinity of solutions in undercompacted shales (higher porosity) should be higher than those in well-compacted ones, providing all other variables are kept constant (p. 25); (2) compaction fluids increase in salinity upon upward migration in a thick shale sequence (p. 274); and (3) ion-filtration does not become significant until overburden pressure reaches about 10,000 psi (p. 238). REGIONAL FEATURES OF COMPACTION

Regional studies of diagenesis must consider all mechanical and chemical effects of compaction, in addition to all other detailed petrologic variations both vertically and horizontally. Numerous case histories are available from the literature of which some are presented below. Various authors showed that many of the diagenetic features are directly or indirectly related to

310

K.H. WOLF AND G.V. CHILINGARIAN

depth of burial or pressure, as well as to temperature, geologic age, fluid composition, variations with time, and others, although not always will one find references made to compaction and compaction fluids. From the published literature and from theoretical considerations, the authors have prepared Tables 3-LI to 3-LIII listing the variables to be considered in regional Compaction studies that may be carried out in the future. Only some of the parameters were considered in the earlier investigations. As in most studies, not all factors can be given the same emphasis for economic reasons and because of numerous other limitations, but a check list, such as the one given here, will assist in choosing those of greatest significance. In presenting the summaries of case histories referring to compaction, it was found best to give them in the order based on the data of publication without losing coherence and continuity. Much of the information was obtained by petroleum geologists during investigations of basin evolution, migration of fluids related to oil accumulation, and related problems. It should be pointed out again that eventually the techniques and concepts developed TABLE 3-LI Factors to be considered in regional studies of compaction (1) Regional facies distribution of coarse elastics, pyroclastics, shales, coal and other organic deposits, and evaporites (very little quantitative information on compressibilities is available on mixtures of lithologies, e.g., (a) sand plus various proportions of clay; (b) sand plus various proportions of clay + silt; and (c) pebbles plus various proportions of sand, silt and clay) ( 2 ) Stratigraphy of all sedimentary deposits (3) Paleotopography, e.g., regional dip variations (4) Paleoenvironments of deposition, which, for example, will control the type of interstitial fluids and diagenesis (and most primary features as listed below) (5) Detailed compositional studies (i.e., granulometric and mineralogic) (6)Detailed textural-fabric and paragenetic studies (e.g., matrix-cement-grain proportions): (a) primary features; (b) secondary features - (i) compactional features, (ii) all other features (7) Detailed mass-property studies: (a) density-porositydepth of burial interrelationships; (b) porosity-permeability-granulometric composition interrelationships (controlling factors: (i) primary features, such as packing, lithology, sedimentary structures, and paleotopography, (ii) secondary features, such as cementation, compaction, leachingsolution, and neoformation of minerals by various processes) (8) Fluid distribution in sedimentary basins (9) Temperature distribution in sedimentary basins (10) Vertical and horizontal variations of all variables listed above (11) Laboratory compaction studies of sediment types encountered in the basin, using uniaxial, hydrostatic, and triaxial compaction apparatuses (12) Comparative investigations between experimental laboratory and natural observations

DIAGENESIS OF SANDSTONES AND COMPACTION

31 1

TABLE 3-LII Texture-fabric indicators of compaction (1)Deformation of matrix (if present), e.g., bending of micaceous grains (2) Interpenetration of grains (3) Type of grain contacts (4)Secondary changes in cementation (if present) (5) Textures and fabrics that might indicate direction of fluid migration (this type of studies has been done on limestones and dolomites with large pores filled with secondary material, but not in the case of sandstones and conglomerates; research and laboratory experiments are needed in this area) (6)Solution features, e.g., stylolites (7) Fracturing of grains, e.g., glass shards in pyroclastics ( 8 ) Conversion of tuffs to zeolitic sediments and bentonite, accompanied by textural changes

by the petroleum geologists can also be applied to the study of ore genesis within sedimentary and volcanic piles. Study by Fiichtbauer fichtbauer (1961) used diagenetic changes on the surfaces of quartz grains to reconstruct the history of oil genesis. He found that the quartz overgrowths in the sandstones of oil fields are more common in the peripheral areas of the oil fields than in the oil-impregnated sections. This indicates that aqueous fluids are required to cause precipitation of silica, whereas oil prevents or interrupts this form of diagenesis. Lowry (1956) had pointed out that the volume of water adsorbed on the surfaces of grains is not sufficient to promote and/or continue silica diagenesis, so that movements of aqueous TABLE 3-LIII Mineralogic changes due to progressive diagenesis-burial metamorphism (1)Clay-mineral changes ( 2 ) Devitrification (e.g., volcanic glass, chalcedony, organic opal) (3) Recrystallization ( 4 ) Changes in mineralogy of coarser detritus, e.g., glauconite to limonite as a result of oxidation; feldspar to zeolite; overgrowths on feldspars and quartz, and corrosion and dissolution of ferromagnesian minerals (5) Changes in mineralogy and paragenetic relations indicating variations in pH, Eh, solubility; temperature, and pressure with geologic time: (a) pH indicator: quartz vs. calcite; (b) Eh indicator: Fe2+ vs. Fe3+ contents; (c) temperature and pressure indicators: zeolite minerals and certain>ypes of phyllosilicates (6) “Crystallinity” of illite (see p. 418 for discussion)

K.H. WOLF AND G.V. CHILINGARIAN

312

fluids through the rock are required. In the basin studied by Fiichtbauer, the sandstones exhibited a range of 1-90% quartz overgrowths. He listed the variables that are responsible for the degree or extent of neomorphism of quartz overgrowths, namely, (a) overburden pressure; (b) history of subsidence and tectonic pressure; (c) time available for diagenesis or age of the rock; (d) chemistry of pore fluids and its changes during diagenesis; (e) interference by carbonate cementation and clay content; and (f) influence of oil migration. The “energy” for silica neomorphism is probably supplied by compaction: (a) mechanical compaction, i.e., turning and rearrangement of grains, which leads to tighter packing; and (b) chemical compaction, i.e., dissolution at grain contacts and deposition at points of smaller hydrostatic pressures. Fiichtbauer (1961,p. 169)reasoned that when one studies individual sedimentary units or reservoir rocks within an oil field, i.e., specimens that were exposed originally to the same pore fluids and have the same age, the same overburden pressure, and the same subsidence history, then the systematic variations in quartz overgrowths of carbonate-cemented and clay-poor sandstones must be a reflection of oil migration. Inasmuch as the presence of oil prevents diagenesis, one must investigate stratigraphic sections that remained unaffected by oil movements to obtain some idea of the type and degree of secondary changes where interruption was at a minimum. Some of the results obtained by fichtbauer are outlined as follows: (1)As to carbonate cementation, Fig. 3-195illustrates that the early diaOO /

OIO

CARBONATE OUARTZ OVERGROWTH MEDIAN

A SP

RESISTIVITY

Fig. 3-195. Restriction of the secondary quartz overgrowths to two dolomitic sandstone lenses in contrast to the remaining water-filled sandstones, borehole Bokel 8, Dogger-fl sandstone “01 ”. (After Fiichtbauer, 1961, fig. 2, p. 170; courtesy Erdol Kohle.)

DIAGENESIS OF SANDSTONES AND COMPACTION YO

OUARTZ OVERGAOWTH

SP

10mV

YO

RESlSTlVlTY QUARTZ MRGROWTH

31 3

SP

RESISTIVITY

210 m

2x

t ”v’ _ p u 29

CARBONATE

KT Nuclei ramoininp withoil

2K m

N0.7

m

No.15

Fig. 3-196. Filtration effect (see text) and coincidence of the present-day with the previous water boundary (= “jump” in diagenesis in the borehole Vorhop-Knesebeck 15). Boreholes Vorhop-Knesebeck 7 and 15; Dogger-0 Sandstone “Ol ” (in both boreholes with water) and “U” (in borehole 7 with water; in the upper part of borehole 1 5 with oil). (After Fiichtbauer, 1961, fig. 3, p. 171; courtesy Erdol Kohle.)

genetic, concretionary dolomite precipitation in the Lias and Dogger sandstones eliminated porosity and prevented the penetration of aqueous solutions. In contrast, the carbonate-free sandstones show quartz overgrowths of up to 70% SOz, whereas in the former instance it was only up to 5%. Similar relationships are exhibited in Fig. 3-196(left side). In contrast, the data presented in Fig. 3-197indicate much earlier carbonate precipitation. (2) To demonstrate a so-called “filtration effect”, i.e., the effect of upward-moving compaction solutions trapped underneath clay-rich layers, Fiichtbauer (1961,p. 170)presented the following discussion. In Fig. 3-196, low at 01,the sandstone shows a 20% quartz overgrowth, whereas in the upper parts of this sandstone it approaches 40%. A similar effect is shown in Fig. 3-196 (right) in the oil-impregnated part. The relationship has been observed in many other cases where the uppermost 1 to 2 m of a sandstone has a higher percentage of silica overgrowths. The parallel increase in staurolite content in Fig. 3-197suggests that in this particular unit no diagenetic corrosion or dissolution has occurred and that silica was precipitated. The silica must have come from an outside source, probablf from upward moving

K.H. WOLF AND G.V. CHILINGARIAN

314

@

SP

@ O/~OUARTZ OVERGROWTH

V~CAR~ONATE

@ yoSTAUROLITE

Fig. 3-197. Tectonic change (inclination) of a structure after oil invasion. (The depths in the diagram are related to the earth’s surface in the structure section NN; also see text for additional explanation.) Dagger$ Oil Field Wesendorf-South, Germany. Numbers refer to well numbers. (After Fiichtbauer, 1961, fig. 4, p. 171; courtesy ErdOJ Kohle.)

compaction fluids derived from deeper clayey units and, possibly, older sandstones undergoing pressure solution. The upward-moving solutions precipitated the silica beneath the impermeable clayey units. Similar mechanisms can be offered to explain carbonate precipitates underneath shales or mudstones. According to Fuchtbauer (1961, p. 170),it is certain that this “filtration effect” was operative prior to oil migration, or was at least penecontemporaneous, because later the remaining water was present in the sandstones only as adsorbed films. An independent explanation also confirms that this phenomenon is an early diagenetic one, i.e., the more pronounced it is, the more intense compaction was or the more compaction fluids were available. Theoretically, the greatest degree of diagenesis as a result of the filtration effect should be present at the upper sections of a thick unit composed of sediments that accumulated quickly and with little or no interruption. With increasing overburden, i.e., with increasing depth in the basin, the filtration effect decreases. Filtration effect is at a minimum in sediments that transgress over older, already compacted or consolidated sediments, because the fluids from the latter have already been pressed out. The effects of filtrating fluids are best observable in deposits which became oil impreg-

DIAGENESIS OF SANDSTONES AND COMPACTION

31 5

nated soon after the formation of early diagenetic features. It should be pointed out, however, that in older sandstones that became water-filled after the filtrating fluids led to quartz diagenesis, the youngest diagenetic precipitation may also have formed quartz overgrowths and thus obliterated all previous diagenetic features. As to the diagenetic differences related to oil migration, Fuchtbauer (pp. 170 and 172) discussed (1)differences within one particular field and (2) differences between oil fields within one large petroleum province. In the case of differences within one oil field, the boundary of the diagenetic effects coincides with the present oil-water boundary. An example of this is shown in Fig. 3-196,where the oil-containing unit has 15-3076 quartz overgrowths, whereas the water-containing part has 40-60% of quartz overgrowths, because diagenesis remained uninterrupted in the latter instance. In between is a transitional unit with specks of oil preserved in it. In addition, the boundary separating the units having higher degrees of diagenesis from those of lower degree does not coincide with the present oil-water contact. Three explanations are possible: (a) oil migration occurred in several steps; (b) the “oil cap” or oil zone was expanded through compaction as a result of reduction in thickness of the oil-bearing unit; (c) the geologic structure was tectonically altered after oil has moved in. Case (a) can be distinguished from the others only very seldom. Early oil migration is probable when a very sharp diagenetic boundary is present within the oil-impregnated unit, but is absent at the present time at the oil-water contact zone. Case (b) is probably applicable in the case presented in Fig. 3-198,where there is an increase in quartz diagenesis towards the bottom. The extension or spreading of the oil cap as a result of compaction is applicable to fields with early oil invasion and impregnation according to the following calculation. Underneath an area of 1m20f the Eldingen Sandstone there were at the time of oil movement about 17.5 m3 of pore space (assuming 50 m thickness and 35% porosity). Since then, the porosity was reduced from 35% to 27%. At the same time, the thickness was reduced at the expense of the pore space by about 5.5 m, from 50 to 44.5 m. Consequently, beneath each 1m 2 of area, the pore space was reduced from 17.5 m 3 to 12 m3. The oil zone must have expanded almost by 50%, therefore. Wherever the sudden change in degree of diagenesis in individual boreholes in an oil field is found at different depths, one has to assume that the structure underwent modification through tilting or disturbances after the oil had migrated into the reservoir rock (case c). An example of this is to be found in Fig. 3-197.The map shows the location of the boreholes 2, 4,5,and 8;borehole 3a is situated far beyond the map. The change in degree of diagenesis in the boreholes 4, 5, and 8 is indicated by diagonally lined areas on the right-hand side. The areas on the left (under the SP curves) indicate presentday position of the bound-

K.H. WOLF AND G.V. CHILINGARIAN

316

t

I

0 2 0 4 0 8 0 1 0 1 0 0

OUARTZ OVERGROWTH,%

Fig. 3-198. Spreading of an oil top through compaction of the sandstone. Borehole Eidingen 55, Liasiu-2 Sandstone. (After Fuchtbauer, 1961, fig. 5, p. 171; courtesy Erdol KOhle. )

ary between oil and water. To reconstruct the form of the structure at the time of silica diagenesis, one has to move the boreholes vertically so that the points at which there is a change in degree of diagenesis form a horizontal line. The distances from the upper surface of the formation to this surface then enables one to reconstruct the former structure. These distances (highs) in the Wesendorf boreholes are as follows: No. 8 = 9.5 m, No. 5 = 8.5 m, No. 4 = 7.1 m, No. 2 < 0 m, and No. 3a < 0 m. One can observe that although the former structure was similar, it was distinctly flatter than the present-day one. Whereas today the borehole No.8 lies 15.7 m above the borehole No. 4, the original structural height difference was only 2.4 m. Because of the high diagenetic quartz content in boreholes 2 and 3a, the sandstones were probably never filled with oil here. In examining the differences between different oil fields within a particular oil province, Fiichtbauer (1972,p. 172) concluded that if two oil fields in an area showdifferences in degree of diagenesis, the oil migration or invasion is older in one case than in the other. An example of this is given in Fig. 3-199,where the Vorhop structure is older than the Hankensbuttel as reflected by the differences in degree of silica diagenesis. Pressure solutionquartz cementation due to the overburden pressure was interrupted in one case by the invasion of petroleum, whereas in the other case the aqueous solutions, which were kept saturated with silica as a result of the compaction processes, caused more extensive silicification. Two examples of the effects of the silica diagenesis on porosity and permeability of the sandstones are presented in Figs. 3-200 and 3-201. In the upper diagrams, the porosity was plotted versus the degree of quartz diagenesis. The porosity magnitude depends on the history of the reservoir rock. The porosity does not show a definite relationship to degree of diagenesis: up t o about 50% of quartz overgrowths, the porosity changes little or not at

DIAGENESIS OF SANDSTONES AND COMPACTION JEDIAN

RESISTIVITY

317

RESISTIVITY

n

&’ 4/20

I I

01

Fig. 3-199.Oil invasions of different ages (older to the left). Boreholes Vorhop 25 (left) and Hankensbuttel M 1 (right). Dogger-fl Sandstone ‘‘01” and “U1”.(After Fuchtbauer, 1961, fig. 6, p. 172;courtesy Erdol Kohle.) At the lower left side, the circle indicates a zone totally impregnated with oil; resistivity curves of left portion of the graph =laterolog 10/100/1,000; resistivity curves a t the left side of figure = ESkl Normal.

all, and only when quartz overgrowth exceeds that value does porosity decrease distinctly (see especially Fig. 3-201).The sediments became compacted down to a porosity of 27-29% without quartz overgrowth development, even if they were oil impregnated. It seems then that quartz cannot dissolve at pressure points, because the silica cannot dissolve in oil. If silica dissolves in the adsorbed water on the quartz grains, it would precipitate in the immediate vicinity. Pressure-solution phenomenon is also unimportant during compaction down to 27-29% porosity, in the case of water-saturated sandstones, because the specimens that have more than 27-295’6 porosity indicate that quartz diagenesis has no effect on the porosity. One would expect an influence, however, in case of pressure-solution occurrence and where the centers of the grains moved closer to each other resulting in pore space reduction. In Figs. 3-200and 3-201,the bottom diagrams illustrate the relationships between permeability and degree of silica diagenesis. Here also, one can observe a-definite intenelationship only when the amount of quartz overgrowths exceeds 50%.

K.H. WOLF AND G.V. CHILINGARIAN

318

0

-

50

I00 0

QUARTZ OVERGROWTH,%

Fig. 3-200

-

50

IW

Fig.3 -201

Fig. 3-200.Dependency of porosity (above) and permeability (below) on the amount of secondary quartz growth. Median diameter of the sample = 0.10-0.16 mm. Eldingen Oil Field, Lias-a-2 Sandstone (-1450-1620 m below the surface). (After Fuchtbauer, 1961, fig. 7,p. 172;courtesy Erdol Kohle.) Fig. 3-201.Dependency of porosity (above) and permeability (below) on the amount of secondary quartz growth. Median diameter of the sample = 0.15-0.24 mm. The low carbonate content was here added to the pore space (see text). Wesendorf-South Oil Field, Dogger-/3 Sandstone (-1480 to 1540 m below the surface). (After Fuchtbauer, 1961, fig. 8,p. 172;courtesy Erdol Kohle.)

Study by Philipp and others Philipp et al. (1963)discussed several new methods in the interpretation of the history of oil migration in their study of the Gifhorn Basin, Germany, containing Middle Jurassic sediments. The results of their investigation can be summarized as follows: (1)A detailed analysis of the structural history, i.e., the trough subsidence and uplift, in space and time yielded information on the age of the reservoir rock development (Fig. 3-202);The thickness of the subsequently-eroded sediments was extrapolated from t6e regional isopach maps (hatched lines, Fig. 3-202). (2)The use of porosity of the shales as a maximum depth indicator is shown in Figs. 3-203and 3-204.A master diagram was employed first considering samples which have never been buried deeper than the present-day

3Im+9MIl m f t

Fig. 3-202. Relationship among age, depth of burial, and accumulation of sediments in the Gifhorn Basin, Germany, based on quartz diagenesis. Location: Liiben to the left and Meerdorf to the right. (After Philipp et al., 1963, fig. 2; courtesy 6th World Pet. Congr., Franfurt/Main.) (See Fig. 3-205.)

K.H. WOLF AND G.V. CHILINGARIAN

320 KAOLIN/CHLORITE RATIO

1111111

INTERSECTION VELOCITY OF CLAYS, m/scc (Liorsic ondDoqger)

‘ “ 1

-

0

POROSITY, %

QUARTZ CONTENT,%

POROSITY, Y.

Fig. 3-203.Variation in diagenetic alterations with increasing depth of burial. Kaolinite/ chlorite ratio has been calculated quantitatively using X-ray diffractograms, whereas the quartz content was estimated from the diffractograms. (After Philipp et al., 1963,fig. 3; courtesy 6th World Pet. Congr., Frankfurt/Main.) Fig. 3-204.Relationship among depth of burial, clay porosity (short vertical dashes), porosity of Tertiary shales, and intersection velocity obtained from sonic logs. The maximum depth of burial (shown by arrows) o r the amount of uplift in Calberlah and Dannenbiittel, Germany, has been determined from the clay porosity and intersection velocity, respectively. The porosity of Tertiary shales is shown by a solid curve. (After Philipp et al., 1963,fig. 4;courtesy 6th World Pet. Congr., Frankfurt/Main.)

depth (Fig. 3-204).X-ray analyses of the shales may support the findings on burial depth, because the diagenetic changes are controlled by the maximum depth of burial as was the porosity, i.e., the kaolinite/chlorite ratios decrease with increasing depth, whereas the quartz content increases (Fig. 3-203). (3)Sonic-log readings were also used as maximum depth indicators (Fig. 3-204),i.e., the interval velocities were employed instead of the porosity values. This method is less time-consuming than porosity determination of the shales and does not require core sampling; however, the exact petrology must be known for proper interpretation of sonic logs. (4) The methods mentioned above are helpful in evaluating the subsidence history of a basin, whereas quartz diagenesis (Fig. 3-205)allows dating of the oil invasion into the reservoir rocks. Although gradual precipitation of quartz cement in clean quartz sandstones is produced mainly by pressure solution without an outside silica supply, some silica precipitated as a result of solubility reduction caused by an increase in salinity of the pore fluids with increasing geologic age and depth of burial. The main controlling variables

DIAGENESIS OF SANDSTONES AND COMPACTION I

M

I

20

LO

30 O/o

50

321 60

quartz grains with overgrowths

W!

11

; I

.25m

m

E

r.ndrtonrr oil-fillrd o w8trr-fillrd A water-fillrd but previously oil-fillrd

Fig. 3-205.Relationship between maximum depth of burial and degree of quartz diagenesis. (After Philipp et al., 1963,fig. 5;courtesy 6th World Pet. Congr., Frankfurt/Main.) The dots are plotted at the maximum depth of burial before oil accumulation; the circles and triangles represent the maximum depth of burial ever reached. Interrupted circles are based o n previous geological interpretations (cf. text). Abbreviations in Figs. 3-205, 3-207 and 3-208:Bk = Bokel; Bo = Bodenteich; Br = Broitzem 4; Ca = Calberlah; D a = Dannenbuttel; Es = Essenrode; G N = Gifhorn-N; GrO = Gross-Oesingen 2 ; Ha = Hankensbuttel-Mitte (in Fig. 3-207 also -N and -0); Ha-Oil = Hankensbuttel-S; Har = Hardesse; HW Hankensbuttel-West; Hz = pit near Bad Harzbuq; II = Ilkerbruch; Lii = Luben; LW = Luben-West; Me = Meerdorf; Oh = Ohrdorf 2; 0 s = Orrel-S 1001; Rii-hi = Ruhme, structural high; Ru-lo = Ruhme, structural low; Ru-HzO = Ruhme, water-filled reservoirs; RU, cse = Ruhme, coarse-grained samples; RU, fin = Ruhme, fine-grained samples; S t = Steinkamp; Th = Thurau 1; Vk = Vorhop-Knesebeck; VN = Vorhop-Nord; Vo = Vorhop; VoH = Vorhop H 1; We = Wesendorf; WN = Wesendorf-Nord; WS = Wesendorf-South; WiS = Wittingen-South 1. The well number appears after the abbreviations, whereas the number of samples is shown below them.

are pressure due t o the overburden and length of burial time. As long as silicification is not too extensive, its degree can be measured very easily under a microscope in transmitted light; however, as more grains merge due to quartz cementation, fewer quartz facets are observable. Philipp et al. (1963,p. 461) defined quartz diagenesis as the percentage of: (1)quartz grains with more than 50% of their surfaces being covered with euhedrd '

K.H. WOLF AND G.V. CHILINGARIAN

322

overgrowths, plus (2)half of the quartz grains with minor silica precipitation. In Fig. 3-205,the open circles represent water-saturated sandstones which have never been oil-saturated during the geologic history. Their degree of quartz cementation has been plotted against the maximum burial depth, which is known in many cases. The distribution of these circles indicates that there is a relationship between quartz diagenesis and depth of burial. The full circles or dots represent oil fields which apparently were formed as a result of early oil migration. To adjust the data, these full circles were plotted versus the maximum depth of burial prior to oil accumulation, as indicated by the quartz diagenesis which has been interrupted by the oil invasion - an assumption proved by the information presented in Fig. 3-205. Based on these well-established test cases, which allow an extrapolation to other similar geologic situations, one can reconstruct more uncertain cases by means of studying quartz diagenesis. The solid triangles inside in Fig. 3-205have a low degree of quartz precipitation, suggesting that the present water-filled sandstones have been oil-filled at an earlier stage. Other geological evidence has confirmed this in some instances. Where the geological evidence for maximum depth of burial was uncertain (vertical lines give a probable range), the solid triangles represent the most likely depth. In stratigraphic traps formed by regression-transgression (e.g., Luben; Fig. 3-195),the quartz precipitation was controlled by the maximum depth of burial prior to the time of transgression. If the oil migration occurred during renewed subsidence after the transgression, but before the rocks reached the previous maximum depth of burial, quartz diagenesis cannot supply information on the time of oil invasion. The broken circles represent the oil pools for which the depth at the time of oil invasion was estimated by using geologic evidence only. The study of quartz diagenesis, however, indicated later oil invasion (full circles or dots) as pointed out by the arrows associated with the broken circles and supported by more detailed geological interpretations. In regard to the scattering of the points, Philipp et al. (1963,p. 463) attributed this t o the different length of burial time. For example, at the Wesendorf-S oil field the maximum depth of burial (1800 m) in the watersaturated wells (WSin Fig. 3-205)is similar to the maximum depth (1500m) before movement of the oil in the oil-saturated WS-wells owing to the late origin of the pool. The water-saturated samples, however, contain 58% quartz grains with overgrowths, whereas in the case of the oil-saturated samples only 34% of the grains have quartz overgrowths. This can be explained by the fact that the water-filled sandstones were submerged to a depth of more than 1200 m for 85 lo6 years until the oil accumulated, whereas the oil-filled ones were buried only for 30 lo6 years. Hence, the different burial time gave rise to different degree of silicification at the same burial depth. Figure 3-205 is based mainly on the wells with the water-

-

-

DIAGENESIS OF SANDSTONES AND COMPACTION

323

.:...D ogger p- sandstones Meerdorf L

I-LL v

20 carbonatemok:-

Fig. 3-206. Influence of carbonate and clay contents on quartz diagenesis. The samples of the upper diagram are poor in carbonate content; the samples of the lower diagram are poor in clay. (After Philipp et al., 1963, fig. 6; courtesy 6th World Pet. Congr., Frankfurtlhlain.)

saturated and the oil-saturated cores, corresponding t o the highest and lowest degrees of diagenesis, respectively. It should be realized that the presence of early carbonate cementation, as well as clay neoformation, could hinder quartz precipitation (Fig. 3-206). (5)Philipp et al. (1963,p. 464) also investigated the intrastratal solution of heavy minerals. Kyanite, staurolite and garnet have been destroyed and, in several instances, a change of kyanite into mica and of staurolite into quartz has been noticed. Figure 3-207 shows that the content of unstable heavy minerals decreases as the degree of quartz diagenesis increases (black symbols). Assuming a uniform chemical milieu, the chemistry of intrastratal solution depends on the maximum burial depth, which, as discussed previously, also affects quartz precipitation. Near the upper and lower sandstone-shale contacts, however, the amount of quartz cement is higher as a result of greater precipitation from compaction fluids, so that here the heavy minerals were not able to be corroded and/or removed (symbols in brackets in Fig. 3-207).It seems that when fluids moving through thick clayey sequences increase in salinity and the salts from these supersaturated compaction fluids may precipitate out immediately upon entering the sandstone. The white or open symbols in the lower part of the diagram represent samples close to the surface, where intrastratal dissolution of the heavy

I

3 60%

I

I

%unst&

-

DSt20 1 DW*rSb 1

OVB

b0,h

D Hz 0I

D

E

4S 1I'0

h.m.

bJ1 25 20 1

matab havY 30% minrmlr

Fig. 3-207. Intrastratal solution of heavy minerals and quartz diagenesis, The percentages of unstable heavy minerals (kyanite + staurolite + garnet) are means of the fractions <0.06, 0.06-0.09, and BO.09 mm. (Signs, see text; for abbreviations, see Fig. 3-205.) (After Philipp et al., 1963. fig. 7; courtesy 6th World Pet. Congr., Frankfurt/Main.)

minerals obviously took place without occurrence of quartz diagenesis. The coarser-grained unstable heavy minerals are the original grains, whereas the fine ones have been modified by dissolution. (6) A very important question that should be answered is: can porosity be used as an indicator of the depth of burial? The data plotted in Fig. 3-208is based on an average of several porosity measurements. Grain size does not appear to have any effect on porosity. In the lower left diagram of Fig. 3-208 the porosity is plotted versus the present-day depth. The lack of distinct correlation is probably due to later uplift and oil invasion (see dis-

DIAGENESIS OF SANDSTONES AND COMPACTION

325

r-POROSITY, O/o

1000

-

E

i In W

0

Zoo0 -

3000 -

k. 0

zoo0

xMo 15

20

25

M

15

20

25

M

E Y

0

9

'

POROSITY, /o ' Oil-filled sondetone8 0 Water-filled sandstone8 \ Present depth - Maaimum depth of burial / Maximum depth of burial before oil accumulalion

Fig. 3-208. Relationship between porosity and depth of burial for Aalenian sandstones. The dots are plotted at the maximum depth of burial before oil accumulation, whereas the circles represent the maximum depth of burial ever reached. For the means presented, only samples containing <4% carbonate, <5% clay (<20 p ) , and 6 2 % by volume of pyrite from wells with a small scattering range of the porosity values have been used. Abbreviations are explained in Fig. 3-205. Lower left-hand diagram: porosities (from the upper graph) are plotted against the present depth. Lower right-hand diagram: porosities of water-saturated sandstones only (from the upper graph), plotted versus the maximum depth of burial. (After Philipp et al., 1963, fig. 8; courtesy 6th World Pet. Congr., Frankfurt/Main.)

cussion under 4 above). The lower right-hand line in Fig. 3-208,which is based only on the "water-filledsandstones, illustrates a definite relationship between the porosity and depth of burial and can thus be used to determine

326

K.H. WOLF AND G.V. CHILINGARIAN

the maximum depth of burial of these water-filled sandstones. The possible influence of oil invasion also has been considered. The porosities of the oil-saturated sandstones (full dots) have been plotted at the maximum depth of burial prior to oil invasion, by making a preliminary assumption, yet to be tested, that the porosity decrease with depth has been interrupted by the oil impregnation. The sandstones filled with oil at shallower depth of about 5 0 0 m and having a high porosity have been compacted from about 35% (water-filled Eocene sandstones have 35% porosity, see 0 s in Fig. 3-208) to 27-31% porosity during further subsidence, in spite of the presence of the oil (Vk, We, Lii, Ha in Fig. 3-208). The compaction is mainly a result of mechanical rearrangement of the grains. As Philipp et al. (1963, p. 465) mentioned, mechanical rearrangement of the grains in laboratory experiments resulted in as little as 27.7% porosity. Although no reservoir rock measurements were made to show the influence of subsidence on the porosity of oil-filled sandstones between a depth of 1000 and 2000 my measurements were available from greater depths (Hur and Me in Fig. 3-208). Based on studies of quartz diagenesis, oil accumulation occurred at a depth of at least 1700 m where the porosity was about 25%. The measured present-day porosity, however, is only about 19% (Fig. 3-208), which may be the result of precipitation from connate waters. From the above data, Philipp et al. concluded that the porosity of sandstones cannot be used at the present state of available information to determine the burial depth during oil invasion. On the other hand, porosities of water-filled sandstones may enable determination of the maximum depth of burial. (7) The variation of properties of hydrocarbons (both gas and oil) with depth have also been discussed by Philipp et al. (p. 466), as they are partly related t o the porosity of the overlying shales and the compaction history in general. Philipp et al. (p. 471) considered the history of basin subsidence and migration of hydrocarbons in individual oil fields based on geological criteria (Fig. 3-195) and quartz diagenesis (Fig. 3-198). They pointed out that the structures presented in their paper appear t o be more gentle than they actually were at various earlier geologic stages because of subsequent compaction. They also mentioned that rates of sedimentation in the central parts of the trough were four times faster than in the peripheral portions. This must have resulted in differential compaction and preferentially directed fluid movements. Also, these investigators pointed out that in certain cases, compaction of the source rocks was interrupted by uplift and erosion, but resumed upon renewed sediment accumulation. Thus, compaction, compaction fluid movements, and related diagenesis can have a complex history and can even be cyclical. It may be of interest that Habicht (1963) in the discussion on the paper by Philipp et al. offered another explanation that is of importance in the study of compaction. He proposed that the principal cause for oil and

DIAGENESIS OF SANDSTONES AND COMPACTION

327

gas migration was not so much the expulsion of oil and gas from progressively compacting sediments, but the increase of temperature in the subsurface. Habicht presented a diagram showing the relationship between quartz precipitation and the time of oil accumulation. The change of depth of burial of the source rock during geological time was also presented. Jankowsky (1963)also employed quartz cementation as an indicator for the relative age of oil versus water. He stated (p. 458)that the formation of petroleum deposits depends on: (a) oil mobilization, which takes place during compaction and is related t o the increase of temperature and pressure in sedimentary basins, and (b) oil migration, first occurring more or less vertically with the compaction fluids and then more or less horizontally into a reservoir rock to form an oil accumulation in a closed structure. If subsequent tectonic modifications occur, e.g., tilting as shown in Fig. 3-209,then remobilization of the fluids can take place. The quartz-diagenesis studies were applied in the latter case. Figure 3-209indicates that the oil trapped by faulting moved from SE to NW within the same stratigraphic or lithologic unit as a result of the tilting. The study of quartz diagenesis of the Dogger Sandstone helped in the reconstruction of these events. In borehole A, the Dogger Sandstone is filled with water today but exhibits little quartz cementation, which indicates pre-Albian presence of oil that prevented more extensive quartz precipitation. There is an extensive quartz diagenesis in borehole B, especially in the deeper sandstone units, suggesting continuous presence of water and presence of small amounts or absence of oil. The Dogger Sandstone in borehole C,which is filled with oil today, exhibits intermediate I.

2'

A

0

C

I

I

I I

I

A BASE OF

'ALB

LOW

n.~i**

of

Hlom

ausm n~~(.n..ii

I

C

I"lOrn*dI.t,

Fig. 3-209. Structural sections to explain the quartz diagenesis. 1 = present-day structure; 2 = pre-Albian structure; 3"=quartz diagenesis in the aquifer. (After Jankowsky, 1963, fig. 6; courtesy 2.Dtsch. Geol. Ges.)

K.H. WOLF AND G.V. CHILINGARIAN

328

CLAY CONTENT,

Oh

Fig. 3-210.Porosity as controlled by clay content. Numbers next to the symbols indicate the sampling depths in cm below the ocean floor. Samples without a number represent surface samples. The curve averages the range of clay contents of the uppermost meter of sediments. I n the diagram in the lower right-hand corner are values obtained from the harbor bay indicating the deviations of porosity from the above-mentioned curve. In Figs. 3-210, 3-211, and 3-212,only samples were utilized that were obtained with a large push-cylinder. (After Fiichtbauer and Reineck, 1963, fig. 4, p. 299;courtesy Sedimentology.) I = sand flat (bank); 2 = mixed (mud + sand) flat; 3 = mud flat; 4 = “surf flat” (= flat with surf and/or breakers); 6 = “flat sea”; 6 = harbor bay or harbor inlet; A = deviation from the upper curve (= average porosity of the uppermost meter) in % porosity; B = depth (in m ) below the upper surface of the sediment.

degree of quartz diagenesis. Together with the information gained from borehole A, this suggests that the oil moved into this sandstone during postAlbian time. Fiichtbauer and Reineck (1963) undertook a study to determine whether or not: (a) sediments of the same grain size, but formed in different depositional environments, have the same porosity, and (b) their condensation ratio* is different. Although the results are related to the material presented in the section on density and compressibility, they are given here because they show some regional or environmental variation and are, therefore, of interest in large-scale compaction investigations. Figure 3-210 shows the increase of porosity with increasing clay content. In pure sands, most porosity values are around 40%. Clays from a young harbor (= Hafenbucht), where

* The condensation ratio = 4w-4 x kv

- $d and &

100

(3-11)

where 6 = natural porosity, and @d = porosities representing the loosest and closest packing arrangements of grains under water, respectively.

DIAGENESIS OF SANDSTONES AND COMPACTION

329

TABLE 3-LIV Summary of average values of densities of water-saturated sediments (after Fuchtbauer and Reineck, 1963, table 11, p. 300) Type of sediments

Density (g/cm3)

Number of samples used

Marsh Dry beach Wet beach Mud-watt Mixed-watt Sand-watt Mud from harbor bay Subaqueous sand bank Shelf

1.427 1.907 1.919 1.538 1.747 1.872 1.332 1.953 1.963

1 15 49 6 7 56 40 7 20

rates of sedimentation are up to 50 cmfyear, have a porosity of about 83% at the surface (determined with picnometer). The clay content of the sediments depends on the rate of sedimentation and other factors of the sedimentary milieu. The clays of the shelf zone (= Flachsee) and Wadden mud (= Schlickwatt) are denser with a maximum of about 70% porosity. Increasing age, coarser grain size, and different sedimentary processes may be responsible for this. The weight of water-saturated sands also depends on the clay content, as presented in Table 3-LIV. As shown in Fig. 3-211, in which porosities were plotted versus median grain size, the three different depositional environments show some variations. These results, however, should not be used for generalization. Noticeable is the small and irregular dependency of porosity on grain size. This, however, is not true when one uses the same sediment samples and performs artificial sedimentation by free fall in the laboratory. In the latter case, there is a distinct decrease in the porosity with an increase in grain size. This cannot be explained by sorting, but may be the result of greater sphericity of the coarse grains that, in turn, results in tighter packing. Similar relationships occur in ancient sandstones, one example of which is presented in Fig. 3-211. These sandstones were buried to a depth of about 1100 m where cementation was prevented by oil impregnation. Fiichtbauer and Reineck concluded that porosity is independent of the grain size in newly-deposited, uncompacted sediments, but interdependence between these two variables increases with increasing burial pressure until there is a normal grain size-versus-porosityrelationship. The increase of porosity with decreasing grain size in recent sands may be compensated by a closer packing arrangement of the finer grains, expressed by a higher condensation ratio. As shown in Fig. 3-212, the condensation ratio increases distinctly with decreasing grain size, which applies to sediments of all depositional environments.

K.H. WOLF AND G.V. CHILINGARIAN

330 I

.

.

.

.

I

.

. . .

I

.

.

.

50%

\ &\

.

h ’

8

ae

i

t

v)

0

a

0

n

\

-30

O!l

012

013

03 MEDIAN DIAMETER,rnrn

OIZ

d3

Fig. 3-211.Relationship between porosity and median diameter in sands. The spread of the values for natural sediments is represented in the area which includes 68% of the values (B). The same area was also duplicated in the right-hand figure, so that a comparison is possible with the artificially accumulated sediments in the laboratory. The diagram in the left-hand corner with a different ordinate axis shows porosity distribution in a Lower Cretaceous sandstone (= “Unterkreidesandstein”). A. Subaerially loosely packed sands as a result of vertical accumulation (dropping down without further shaking). B. Recent sands. C. Subaqueously loosely packed sands as a result of vertical accumulation. D. Subaqueously moved sands with tighter packing. 1 = surf beach o r shoreline beach; 2 = sand flat; 3 = flat sea. (After Fiichtbauer and Reineck, 1963, fig. 5, p. 301;courtesy Sedimentoiogy. )

One should notice, however, that the organically-reworked sediments (by boring and burrowing) have a higher condensation ratio. The results presented in Fig. 3-212are in agreement with those obtained by civil engineers, i.e., that the high energy required to obtain maximum values of condensation of coarse sands is usually not available in natural sedimentary environments, in contrast to the lower energy required for finer sediments. Only during compaction will the coarse sediments be “condensed”. As shown in Fig. 3-210, the porosity of the upper layers of quickly-accumulated clays is 83%, whereas that of slowly deposited clays is about 70%. The decrease of porosity from about 83 to 76% (a difference of about 7%) in the upper 4 m of

DIAGENESIS OF SANDSTONES AND COMPACTION

lo 90-

2

331

I

80 -

70-

'z 0

5

m

60-

50-

W

n

z

8 LO30200 O \0" 0

10-

I

0.2

0.1

0.3

MEDIAN DIAMETER, mrn

Fig. 3-212. Relationship between "condensation" and median diameter (see text). 1 = sand flat; 2 = surf flat (= flat with surf and/or breakers); 3 = shallow sea; 4 = shallow sea (well agitated); 5 = shallow sea (partly agitated); 6 = shallow sea (strongly or well agitated). POROSITY,% IS

20

2s

"00

35

30

//RECENT /

/

OL/GOCENE D

500

000

/

E

/

/

, ,

//

EOCENE

/

2

5 3

m

0 lL

r a n 2500

,'

3000

I

/

/

/

PURE SANDSTONES

/

"/#MAT

Fig. 3-213. Decrease- of porosity in pure sandstones with increasing maximum burial depth. (After Fuchtbauer and Reineck, 1963, fig. 7, p. 304; courtesy Sedimentology.)

K.H. WOLF AND G.V. CHILINGARIAN

332

sediments is shown in the right-hand corner of Fig. 3-210.The porosity in the upper 1 0 0 m is very variable and may depend on the mineralogical composition as well as transportation-deposition mechanisms. The type of clays present plays a major role in determining porosity of clay deposits, e.g., illite versus montmorillonite. The same arguments apply to sands and sandstones with a clay matrix. Figure 3-213 presents one example where pure sandstones decrease in porosity from about 40% at the surface to about 14% at a depth of 3300 m. Usually, straight-line relationships between porosity and depth of burial occur only where the sediments are matrix- and cementfree.

Study b y Adams Adams (1964)studied a sandstone formation which had undergone extensive diagenetic alterations on a regional scale, that influenced its reservoir properties. This formation of Pennsylvanian age consists of two basic sandstone types (Fig. 3-214): (A) nearshore, clean, relatively well-sorted, nonglauconitic, non-calcareous sandstone deposited under high-energy conditions, and (B) sandstone representing seaward facies of type-A sandstone, which usually is argillaceous, poorly-sorted, chloritic or glauconitic, and calcareous, and was deposited by a medium of lower energy than that required by type A. The diagenetic features include effects due to pressure solution

-+TYPE

'6' SANDSTONE

Fig. 3-214. Major depositional facies relations of typical Lower Morrowan sandstone (Pennsylvanian) Anadarko Basin, northwestern Oklahoma. (After Adams, 1964, fig. 3, p. 1570; courtesy Am. Assoc. Pet. Geologists.)

DIAGENESIS OF SANDSTONES AND COMPACTION

(A) W S T D N E TYPE A FABRIC CLASS I

(8) SANDSTONE TYPE A FABRIC CLASS II

(C) SANDSTONE TYPE A FABRIC CLASSI AI

+FAWN cussmi I

n

333

IE) SANDSTONE TYPE A FABRIC CLASSII[ B I

IF) SANDSTONE TYPE B FABRIC CLASS JE B 1-2

(D1 dmmm oB W

” E

TYPE A

SANDSTONE TYPE B FABRIC CLASS B 1-3

m

SANDSTONE TYPE B FABRIC CLASS a B 3

Fig. 3-215. Thin-section illustrations of major fabric classifications. (After Adams, 1964, fig. 5, pp. 1572-1573; courtesy Am. Assoc. Pet. Geologists.) For description of the classes, see Table 3-LV. C is porous.

and presence of several cements, i.e., quartz, calcite, dolomite, and clay. Replacements, decementation, and corrosion are common. Adams discussed: (1)pressure solution; (2) quartz overgrowth; (3) stylolites; (4) sutured, concavo-convex, and other types of contacts versus floating grains; (5) influence of clay on pressure solution; (6) cementation related to compaction; (7) distribution of calcite versus dolomite cement; (8) clay matrix and cement distribution; (9) etched quartz grains; (10) porosity (original and secondary) distribution prior and subsequent t o mechanical and chemical diagenesis. He found a regular and predictable distribution of many of these features from sandstone facies of type A t o type B. Adams offered a petrofabric sandstone classification scheme in which subdivisions are based on degree of pressure solution with various modifiers, including grain outlines and contacts, amount and type of cement(s), and porosity (Table 3-LV) (see modifications made by Pettijohn et al., 1972). Sketches of each of these major petrofabric types are presented in Fig. 3-215. According to Adams, “diagenetic facies”

DOWNDIP DEPOSITIONALLY

STEP 3 Well No. 2

VALUATION: Wildcot well no. I hos encountered good 90%poy. No clue to optimum dirrction for field development. ECOMMENDATION: Study rocks, clossify ond integmta with rock geometry.

EVALUATION: Well no.2 encountered Sondsto$e,yB," we11 developed and Sandstone A tqht. No Clue to optimum dirtction for odditimal development: RECOMMENDATION: Study rocks,classify and integrate with rock geometry.

-

STEP 4 Well No.2

STEP 2 Well No. I

4

13,. SANDSTONE TYPE B FABRIC CLASS&! BI GAS fP/fl, MCFD SANDSTONE TYPE FABRIC CLASS

m

Sample and petrographic anolyses slmw pay to Be 8econdory. Gbomrtry of such a yrervoir is unpredickblr. Sandstone 8 is fine groined with FobricClorsI ond is pmbobly ot updip edge of porous, high energy sondstone. ECOMMENDATION: Drill dopqdip depositionally from Sondstone 8 .

VALUATION:

EVALUATION: Samplr ond pstrogrophic mol ws show pay to be in high ensrgy ronds!me with pood prediction possibilities. RECOMMENDATION: Drill along Cposilionol strike of Sondrtone "8':well no. 2

DIAGRAMMATIC SKETCHES SHOWING APPLICATION TO HYDROCARBON EXPLORATION (Wells No.1 ond 2 Are Appror

1

I Mile Aport

I

MECHANICAL LOG POROSITY

= SANDSTONE TYPE "A"

4I

= SANDSTONE TYPE "8" * SPONTANEOUS POTENTIAL LOG

Fig. 3-216. Sketch of an actual case within study area (subsequent to original study in 1961) depicting importance of detailed rock analyses to hydrocarbon exploration. (After Adams, 1964, fig. 6, p . 1574; courtesy Am. Assoc. Pet. Geologists.)

DIAGENESIS OF SANDSTONES AND COMPACTION

335

TABLE 3-LV Criteria of fabric classification and applicable descriptive terms (modified from Gilbert, 1949,p. 13,and Thomson, 1959,p. 100, in Adams, 1964,table I, p. 1579)

Fabric Class I High degree of pressure solution, flattened grains, no porosity, no cement, sutured contacts Fabric Class I I Moderate degree of pressure solution, approximately equidimensional grains, little or no porosity, little or no cement, primarily concavo--convex contacts with some flat, sutured contacts (A) Cemented by: (1)quartz overgrowths; (2)quartz overgrowths with minor amounts of calcite; (3)quartz overgrowths with some clay Fabric Class 111 Minor degree of pressure solution, mostly original grain outlines, often some porosity, long or tangential contacts (A) Poorly cemented (but normally tightly packed): (1)quartz overgrowths; (2) calcite; (3)dolomite; (4)clay (B) Well cemented: (1)quartz overgrowths; (2)calcite; (3)dolomite; (4)clay Fabric Class IV No pressure solution, original grain outlines, good porosity and/or cementation with no porosity, tangential contacts and floating grains (A) Cemented but porous: (1)calcite; (2) dolomite; (3)clay (B) Cemented, no porosity: (1)calcite; (2) dolomite; (3) clay

could be established based on his classification, and Fig. 3-216represents an actual example of his approach. Although he applied it to oil and gas exploration, eventually such diagenetic investigations will probably be used in solving ore genesis problems, e.g., of uranium ores of the Colorado-Plateau type and copper ores in the red-bed sandstones. Study b y Horn

Horn (1965) undertook a regional study of the reservoir properties of a Jurassic sandstone, i.e., porosity and permeability that depend on lithology and facies types as well as on the characteristics of the pore system (content of matrix and cement, and degree of compaction, for example). The following diagenetic-authigenic minerals were found in pores of the sandstones: chamosite, siderite, pyrite, brookite, quartz, feldspar, kaolinite, barite, ankerite and calcite. Calcite precipitation caused the greatest reduction in porosity from an initial value of 40% to 1.5-7%. The sandstones cemented by kaolinite have 6--11% porosity. The other authigenic minerals did not influ-

336

K.H. WOLF AND G.V. CHILINGARIAN

ence porosity greatly except on a microscale perhaps. Siderite occurs in thin zones associated with siderite concretions. Ankerite, which is confined to one particular sandstone unit, locally reduced the porosity to 5%. Pyrite, brookite, feldspar, and barite have minor influence on porosity, whereas quartz was very important in reducing porosity. In discussing the process of silicification, Horn (p. 251) stated that precipitation of SiOz can be prevented or interrupted by (1)filling of the pore spaces by non-aqueous fluids such as oil, or (2) precipitation of a mineral on the quartz grains so that the latter cannot come in contact with the interstitial fluids and, consequently, no silica can be chemically deposited. Presence of large amounts of matrix in a sandstone also would prevent the entry of fluids. Heath and Dymond (1973) have made similar observations on recent sediments. They stated (p. 181) that on the basis of dissolved silica, the pore waters from North Pacific deep-sea sediments fall into two groups. Pore water of oxidized deposits contain 22-35 ppm, whereas those of reduced sediments contain 10-19 ppm of silica. The above difference in silica concentration remains unexplained. Heath and Dymond (p. 184), however, offered a conceivable possibility, namely, that the fine-grained silicate particles in the oxidized sediments were rapidly coated with ferric hydroxides, which inhibited the attainment of equilibrium; whereas in the reduced accumulations reversible reactions, i.e., decomposition of the silicate minerals, are much more likely to take place (see their equations 1and 2). Horn (1965, p. 252) found that rims of radially-arranged chamosite prevented formation of quartz cement. Chamosite is present only on primary clastic grains and not on any diagenetic precipitates, which indicates its very early precipitation. As Fig. 3-217 indicates, there is development of two types of chamosite, i.e., types A and B. Variety A, which occurs as welldeveloped chamosite rims at least 3-5 p thick, prevented secondary silicification so that the average porosity of sandstone is 22.8%. The type-B chamosite occurs as thin rims (less than 3 p ) , so that precipitation of silica was possible and the average porosity of the sandstone was reduced to 16.2% (Fig. 3-217). The variation in the percentages of chamosite of types A and B occurs both horizontally and vertically within the sandstone units. By examining the regional distribution of chamosite zones, one can explain also the regional variation of porosity. The chamosite formation followed shortly after sedimentation of the clasts. Thus, one can expect some relationship of chamosite occurrence to the sandstone facies, because during the very early diagenetic stages, when the chamosite was formed, there was probably no complete renewal of the intrastratal pore fluids. This assumption is confirmed by: (a) the fact that the rims of chamosite are preferentially present in clay-poor, homogeneous sandstones, but are absent in sandstones with clay laminae, and (b) the association of the chamosite with the chamosite

331

DIAGENESIS OF SANDSTONES AND COMPACTION DiSTRlBUTlON POROSITY

0EPTH.m SEDIMENT

OF

-m &0 a

20 10

15 20 25 POROSITY

10

0

1

W 1

10

100

1000

10

15

20

25

B L

m 5

B

B2 3 + 6

' 7

Fig. 3-217.Dependency of porosity and permeability (averaged values of the measurements normal and parallel to the bedding) on the chamosite content in oil-free, upper coarse-grained "Haupt" Sandstone of the borehole Plon-Ost 105. According to thinsection studies, type A = well-developed chamosite rims, silicification is absent ( 4 ) ;type B = thin chamosite rims, strong silicification (5);transitional between types A and B (6); no thin-section observation (7); fine-grained sandstone (1); mudstone (2); calcitecemented sandstone (3); loss of cores (KV). (After Horn, 1965, fig. 4,p. 251 ; courtesy Erdol Kohle.)

oolites. Figures 3-218A7B give the regional distribution patterns of chamosite; the top diagrams refer to the upper, coarser sandstones (11) and the lower diagrams to the lower, finer sandstone (I) units.'In unit I, there is an elongated SW-NE zone of clay-free sandstone containing conglomerate horizons; an increase in grain size, in general, reflects more turbulent depositional environment. Especially here one finds chamosite oolites and authigenic films of chamosite on the grains. Degree of silicification was low (Fig. 3-218,E) in unit I and high porosity was preserved (Fig. 3-218,F). Clayey sandstones with normal laminations and cross- and flaser-bedding, accumulated outside this elongated zone. Inasmuch as here the rims of chamosite were only poorly developed or absent, a reaction between the fluids and the quartz grains was possible and silicification with consequent porosity reduction took place (Figs. 3-218, D,E, and F). Similar observations have been made in the upper sandstone of unit 11, but the zones of stronger chamo-

K.H. WOLF AND G.V. CHILINGARIAN

338

il

LOCATION

- 1

ow---

MAP

uI.

0 2

TYPES OF SEDIMENTS AND STRUCTURES

I POROSITY, %

TYPE OF CEMENT

1

.A&?

12.8

3 5. 3 6.

20.8

17.3

El?

B 8.

Fig. 3-218.Comparison between sedimentary facies development, cementing material and porosity in two units of the Dogger-/? “Haupt” Sandstone of the Plon-East Oil Field. (In the sketch map to the left, the depth lines of the “Haupt” Sandstone are 100 m apart.) 1 = homogeneous, clay-free sandstone; 2 = clayey sandstone; 3 = unoriented flakes (= plates) of chamosite; 4 = oolites of chamosite; 5 = horizontal thin bedding; 6 = flaser bedding; 7 = cross-bedding; 8 = conglomerate horizons; 9 = authigenic chamosite (bank); 1 0 = extensive secondary silicification. A, B, C = upper coarser-grained unit of ‘‘Hauptsand”; D, E, F = lower finer-grained unit of “Hauptsand”. (After Horn, 1965, fig. 5, p. 252; courtesy Erdol Kohle.)

site development are more pronounced (Figs. 3-218A,B, and C). Very porous sandstones with an average porosity of 21.9-23.5% correspond to type-A chamosite precipitation (Fig. 3-218,C, lower right), whereas the maximum porosity of 18.3%in the upper part of Fig. 3-218,C is present in sandstones having the type-B chamosite genesis that is accompanied by weak or reduced silicification. In his discussion on the influence of oil invasion on diagenetic silica precipitation, Horn (p. 253) stated that inasmuch as the porosity in the sandstones discussed above were determined by the synsedimentary chamosite genesis,

339

DIAGENESIS O F SANDSTONES AND COMPACTION

later differences in the supply of silica were probably less important. On the other hand, in clay-poor sandstones, in which the grain surfaces are not covered by films of chamosite or chlorite, it seems possible that quartz diagenesis was influenced by early oil and gas invasions. These latter conditions apply in another oil field investigated by Horn. As Fig. 3-219,Aillustrates, the fine-grained and coarse-grained sandstones are interbedded with the clayey sandstones, which contain clay laminae. Locally, these sandstones are replaced by sandy claystones (Fig. 3-219,A).In contrast to the earlier example mentioned (Fig. 3-218),Fig. 3-219,Cshows that the lines of equal porosity are parallel to the contour lines of the sandstones, despite the presence of a distinct facies differentiation. This suggests that the porosity distribution is not controlled by primary sedimentary features. The more intense cementation or lithification in the north is the result of more intense silicification (Fig. 3-219,B).Inasmuch as there is no relationship between the LOCATION

MAP

I P E OF SEDIMENT

I , -

TYPE

OF

CEMENT

POROSITY, o /‘

OOO

_.-

21.1

:EL

@L

-

2.

Fig. 3-219. Relationship among sedimentary facies development, type of cement and porosity in the upper coarse-grained “Haupt” Sandstone of the Preetz Oil Field. (In the diagram to the extreme left, the depth lines of the “Haupt” Sandstone (“Hauptsand”) are 100 m apart.) a = strong silicification; b = weak silicification; I = boreholes examined; 2 = facies boundaries; 3 = claystone; 4 = fine-grained sandstone; 5 = coarse-grained sandstone; 6 = claystone with illite-muscovite. (After Horn, 1966, fig. 6, p. 253; courtesy Erdol

Kohle.)

K.H. WOLF AND G.V. CHILINGARIAN

340

lithification and the sedimentary facies, it appears that differences in the cement distribution are the result of diagenetic processes. The absence of quartz precipitation in the south may have been the result of oil invasion that prevented the movement of aqueous solutions. The northern area was situated near the oil pool’s margin, filled with water, so that silica precipitation was not hindered and the sand was cemented more intensely, as exhibited by the presence of 5-6% more cement (Fig. 3-220).The E-W boundary between boreholes 6 and 9, exhibiting striking differences in diagenesis, does not coincide with present-day marginal water boundary, which extends in a E

AVERAGE VALUE OF POROSITY, 10

W J

> W -1

I

2400-

”

“

”

’

PREETZ

15

”

’

”

*

”

O‘/

20

25

12 a

e

0

m

N W

4

2 500-

z 0

2600

A

W

m

a

-

v)

I-

a

3

I

2 700 O A

W

I

I-

LL

r-

2700

P LO N

- o ST

z 3

E

m

2800

m W

0”W

n

2900

W

I I LL

0

3000

rT >

$

3100

8 rT

s

3200

W

f

U

0

I I-

a W n

3300

b. o c.

15

20

AVERAGE VALUE OF POROSITY,

25 O/O

Fig. 3-220.Relationship between porosity and depth of burial of the “Haupt” Sandstone (“Hauptsand”) in the oil fields Plon-East and Preetz. a = oil-free Gamma Sandstone, Preetz 4; b = upper coarse-grained part of “Haupt” Sandstone; c = lower finer-grained part of “Haupt” Sandstone. (After Horn, 1965, fig. 7 , p. 254; courtesy ErdOZ Kohle.)

DIAGENESIS OF SANDSTONES AND COMPACTION

341

N-S direction at about 2630 m depth parallel to the contour lines of the sandstone (Figs. 3-219and 3-220).Oil migrated later into the northern part of the sandstones after the subsurface fluids deposited silica cement. The sandstones in borehole 9 now contain secondary water. Remains of oil and the relatively high porosity indicate that the rocks contained oil earlier. Horn (p. 254) compared the silicification of the sandstones in the PlonOst and Preetz oil fields and found that in the former sandstones the porosity is dependent on the chamosite content, whereas in the latter the silica cementation is controlled by earlier oil invasion. A comparison between these two rock types is desirable because these sandstones are of the same age, occur at the same depth in subsurface, and originally contained similar pore fluids, but exhibit differences in facies and diagenesis. In both cases, there is a linear decrease of the average porosity with increasing depth (Fig. 3-220).The effect of increasing compaction here is related to the pressure solution of the quartz grains at the grain contacts with subsequent or simultaneous precipitation of silica from the saturated solutions at other localities. This mechanism has been described in detail by many investigators in the literature. In the Preetz oil field, silica precipitation took place more or less at the site of pressure solution. In the Plon-Ost oil field, pressure solution also occurred at grain contacts, where the rims of chamosite were missing and are 'cly*>z,?nt only as linings of the pore walls. On the other hand, in the chamosii.e-rich sandstones, the silica precipitation from saturated solutions was hindered. It is possible that the silica was supplied from more remote localities, where pressure solutions due to the overburden pressure was possible and the fluids moved into the clayey sandstones to take part in the quartz diagenesis. The silica may also have been derived from feldspar decomposition and from clays during mineral transformations. Hence, there is a considerable variation in the average values of porosity at the same depth in the sandstones of the Plon-Ost oil field and only a small decrease of the average porosity values of 1-2% per 100 m of depth, in contrast to the gradient of 5% per 100 m reported from the sandstones in the Preetz oil field (gently versus steeply sloping average curve, Fig. 3-220).To what extent the intense decrease in porosity in the Preetz oil field is the result of overburden pressure and whether other factors are involved, has not been determined because a depth interval of 200 m which was studied is not sufficient. A periodic or cyclic oil invasion, for example, may have effects similar to those of overburden pressure, when silicification can occur in lower horizons but is prevented by the oil at higher horizons, until the lower units also receive oil and the silica precipitation is interrupted. As discussed above, high porosities can be maintained in quartzose sandstones when pressure solution is prevented and/or when silica-saturated solu-

342

K.H. WOLF AND G.V. CHILINGARIAN

tions cannot react with the quartz grains because of the presence of rims of minerals around the clastic constituents. As reviewed above, silicification can be prevented also by invasion of oil. Although the chamosite rims can preserve the high porosity to a greater depth, inasmuch as these rims prevent or diminish cementation, the rate of oil migration will be less, because of greater adsorption of oil to the surfaces of the grains. The presence of rims decreases the size of the pores and at the same time increases the surface area per unit of rock volume. The presence of chamosite rims, therefore, is not only important in controlling presence and degree of pressure solution and, consequently, porosity and permeability, but also has to be considered in evaluating the subsurface hydrodynamic flow. The degree of compaction may vary from layer to layer and may depend on the amount of chamosite rims and matrix formed during early diagenesis. (See Table 3-LVI.) TABLE 3-LVI Dependency of the porosity on the chamosite content in sandstones of the same facies (“Haupt” Sandstone Plon-East 102; length of sample = 20 cm; permeability measurements parallel and normal to bedding; (after Horn, 1965, table 1, p. 253) ~

Depth (m)

3239.9 3241.7 3243.6

Chamosite content

4-

-

Secondary silicification

-

+

+

-

-

+

+

-

-

+

~~

~

~~

~~

Porosity (%)

Permeability (md)

Il*

I**

II

1

(%)

22.3 5.9 25.6 9.8 20.8 9.8

18.1 7.5 25.2 9.0 19.5 9.0

125 -

4.3 0.1 163 2.5

1.2 2.1 2.5 0.9

22 1.4

1.3 3.3

189 2.4 83

-

Amount dissolved in acid

* 11 = parallel to bedding; ** 1= perpendicular to bedding. Diagenetic features and chemistry o f pore fluids (Selley, 1966) Selley (1966), in his discussion on diagenetic variations on a regional scale (both vertically and horizontally), suggested that the diagenetic features may be related not only t o pressure solution or burial pressure but also to the chemistry of the pore fluids, for example. In Fig. 3-221, he presented three different units that originated under different environments. Pressure solution and cementation of quartz is confined to those units that accumulated under oxidizing conditions, in contrast to the lacustrine-marine sediments

DIAGENESIS OF SANDSTONES AND COMPACTION

343

Fig. 3-221. Diagram illustrating the relationship among mode of diagenesis, lithofacies, and presumed depositional environment. (After Selley, 1966, fig. 4; courtesy Proc. Geol. Asso c. )

which contain large amounts of chlorite, carbonate, and quartz cements. Cementation does not seem to have been accompanied by pressure solution. The effects of compaction and depth o f burial (Fuchtbauer, 1967) Fuchtbauer (1967a) presented the results of an investigation in which he directly considered effects of compaction and depth of burial. In the section on quartz diagenesis and mechanical compaction, he showed (Fig. 3-222,A) that the porosity of the “Dogger beta” quartz sandstone decreases with increasing depth. The curve nearly coincides with the dashed curve presented by Proshlyakov (1960, in Maxwell, 1964). Examination of the grain surfaces revealed the cause for this porosity decrease, i.e., the percentage of quartz grains, which show secondary overgrowths on crystal faces, increases with depth (Fig. 3-222,B). As the silica is dissolved as a result of pressure solution at points of contact, the grains slip into denser packing (Fig. 3-223,A). The latter figure also shows that the amount of pressure solution required is small. Fiichtbauer (p. 355) pointed out that the rock volume to be dissolved is in general smaller with increasing steepness of the contact faces of the grains. In Fig. 3-223,B, changes in volumes are based on the most unfavorable case of horizontal contact planes. Only 1.5% of the material must be dissolved to result in compaction (up t o 50%), as graphically shown by curve A in Fig. 3-224. The amount of quartz which has to be dissolved evidently increases with compaction, because contacts become longer and the number of horizontal contacts increases. Consequently, as pointed out by Kchtbauer (p. 355), the “mechanical slipping (rearrangement) of grains dominates during early compaction, whereas chemical compaction (pressure solution) is predominant during later stages”. In determining the dissolution of quartz with depth, as shown in Fig. 3-222,B, Fuchtbauer employed curves A and B in Fig. 3-224 and the porosity curve of quartz sandstone in Fig. 3-222,A. The general shape of the curves is probably correct, but the maximum amounts of dissolved quartz may be unrealistic. The dissolution of silica is minor in the upper 1000 m of sediments, but increases roughly

K.H. WOLF and S.V.CHILINGARIAN

344

OUARTZ GRAINS WITH OVERGROWTHS, %

POROSITY, %

0

5

10

DISSOLVED OUARTZ REPRECIPITATED, %

Fig. 3-222. Porosity and Si02 migrations in relation t o depth of burial. The maximum depth of burial is the greatest depth t o which a sandstone was ever buried. (After Fuchtbauer, 1967a, fig. 1,p. 354; courtesy 7th World Pet. Congr.) A. The values for calcareous sandstones were taken from many measurements of the “Bausteinschichten” (= Chattian Molasse; Fiichtbauer, 1964, fig. 22). The curve for the quartz sandstones is for the “Dogger beta” (after Fuchtbauer and Reineck, 1963). H I = average value for silicified sandstones; H2 = sandstones with diagenetic chamosite seams that hindered quartz diagenesis (after Horn, 1965). (Dashed curve is after Proshlyakov, 1960, in Maxwell, 1964.) B. Lower left = maximum amount of quartz dissolved by pressure solution (see text). Upper right = quartz diagenesis, i.e., percentage of quartz grains with secondary overgrowth (after Philipp et al., 1963). Fuchtbauer (1961) defined it as: [100(0.5 b + c ) / ( a + b + c ) ] , where a = number of grains with very few crystal faces, b = grains with some crystal faces, c = grains with crystal faces covering the larger part of the surface. The counting was done with microscope using grains of 0.12 to 0.15 mm fraction on a dry glass slide. These values, however, do not depend on the grain size.

A

B

Fig. 3-223. Mechanical and chemical compaction. A. Incline(- contact planes - mechanical compaction is predominant. The rock volume to be dissolved (black) is even smaller if rotational movement occurs. B. Horizontal contact plane - exclusively chemical compaction. Much more material must be dissolved in this case in order to achieve the same amount of compaction. Its amount is independent of grain size as deduced from geometric considerations. (After Fuchtbauer, 1967a, fig. 2, p. 365; courtesy 7th World Pet. Congr.)

DIAGENESIS OF SANDSTONES AND COMPACTION

COMPACTION, %

345

--c

Fig. 3-224. Theoretical relationship among degree of compaction, porosity (curves B and C), and the amount of rock volume dissolved (curve A ) for the model presented in Fig. 3-223,B. Curve A = dissolved quartz(%), ordinate to the right; curve B = prosity (quartz removed), ordinate to the left; curve C = porosity (quartz reprecipitated), ordinate to the left. (After Fuchtbauer, 1967a, fig. 3, p. 356; courtesy 7th World Pet. Congr.)

linearly with depth; this, in turn, correlates with the linear increase of silica overgrowths below a depth of several hundred meters (Fig. 3-222,B).The upper 1000-1500 m of sediments are, therefore, the domain of mechanical compaction. It seems that both sorting and roundness of grains influenced compaction, roundness being particularly important. The finer-grained sands are more porous than the coarser-grained ones having a comparable clay content, which does not agree with the relationship postulated by Weyl (1959) for the effective range of pressure solution. The appreciable difference in porosity is indicated by the short line marked 0.08 mm in Fig. 3-222,A for some “Dogger beta” rocks. Fine-grained “Bentheimer Sandstein” has a porosity of 27% at a depth of 1100 m, whereas the coarsegrained sandstones have a porosity of 22% (see Von Engelhardt, 1960, figs. 11 and 49). The differences in sorting may be significant, but do not seem to explain completely the higher porosity of the finer-grained sandstones. As Von Engelhardt (p. 21) stated, the higher porosity values of finer sandstones can be explained by the larger number of grain contacts/unit volume of sediment resulting in a higher resistance against compaction. As this has not been found to apply to all petrographically similar fine-grained sediments, other factors must also be active. A more important factor may be the roundness, as it is commonly better developed in comer sandstones (exclud-

K.H. WOLF AND G.V. CHILINGARIAN

346 4

3 2

In

I

B

cn

w z

3

3

2

0 2

0

a

I

3

I

C

2 I

GRAIN SIZE, m m

Fig. 3-225. Rounding (Russell-Taylor) in relation to grain size. Ordinate (roundness): 1 = angular; 2 = subangulir; 3 = subrounded; 4 = rounded. This number was multiplied by the number of grains falling into the respective roundness class. Then the sum of these products was divided by the number of all grains. Each point corresponds to one grainsize fraction. The median diameter is shown by an open circle for each sample. A = “Bentheimer” sandstone from Scheerhorn (Hecht et al., 1962), oil-saturated, 4 samples; B = “Dogger beta” sandstone from Hankenbiittel, oil-saturated, 2 samples; C = “Dogger beta” sandstone from Luben-West, water-filled, 2 samples. (After Fuchtbauer, 1967a, fig. 4, p. 356; courtesy 7th World Pet. Congr.)

ing “textural inversion” of Folk, 1968), and, therefore, the coarser grains “slip” better. From Fig. 3-225, two general relationships become clear: (a) roundness increases with grain size, as has been reported from numerous other studies; however, it reaches a maximum in the coarse sand range (curve A ) ; (b) identical grain fractions are better rounded in the fine-grained sandstones than in the coarse-grained sediments. This can be explained by the differences in transportation mechanisms, namely, rolling versus suspension. The median rounding (= rounding of the median-sized grains, shown by circles in Fig. 3-225) can, therefore, be larger in coarser sandstones (curves A and C) or equal to that of fine sandstones (curve B ) . It follows then, that should rounding affect porosity, the curves A and C should show a relationship between grain size and porosity. But curve B (center of Fig. 3-225) does not show such an interdependency, as demonstrated in Table 3-LVII. According to Fiichtbauer (p. 357), there is an important influence of the degree of rounding on compaction during the mechanical stage of diagenesis resulting in a reduction of porosity. Inasmuch as the specimens contain on the average only 11%quartz grains with overgrowths, the amount of pressure solution must have been comparatively small.

347

DIAGENESIS OF SANDSTONES AND COMPACTION POROSITY ( d o t s )

FILLED

Q U A R T Z OVERGROWTHS

CARBONATE G + S

- 1510 m

,

,

,

f.’L/A5,

I

*

:

,

,

, ,:

j

,

,

I

Md a18 0.20

0.1

1

0.08

0.10

TERTIARY OILFILLED

+ m

Fig. 226. Relationship between the L-gree of quartz diagenesis (centei,, pore filling (left), carbonate content (right), and distance from the sandstoneshale contact for three sandstones of the “Dogger beta” Formation. (After Fuchtbauer, 1967a, fig. 5, p. 357; courtesy 7th World Pet. Congr.) Top = Wesendorf-South 2 (after Fuchtbauer, 1961); center = Hankensbuttel; bottom = Bodenteich; left = median diameter in mm, porosity, and SP curve; right = G + S = garnet + staurolite in the heavy-mineral fraction (after Dmng, 1965).

348

K.H. WOLF AND G.V. CHILINGARIAN

The second stage of compaction is the chemical one, characterized by increased pressure solution combined with precipitation of silica on quartz grains. This can lead to complete silicification of the sandstone as long as no other cementing mineral occurs. Figure 3-226, which shows the source of silica, consists of three parts: (a) top: a section through a water-filled rock; (b) center: a section through an oil-filled rock; and (c) bottom: a section through a formerly oil-filled, but now water-filled sandstone. Each one of these cases is considered separately. (a) The amount of secondary quartz overgrowths is high in the water-filled section, except for portions with carbonate cements. The latter preserved an early stage of diagenesis, which is evidenced by the presence of a number of unstable heavy minerals (e.g., garnet and staurolite). (b) Less quartz precipitation occurred in oil-impregnated sandstones. The minor amounts of quartz present apparently have been preserved by the oil. Using the upper right-hand side (hatched area) of Fig. 3-222, one can deduce that the depth at the time of oil migration-and impregnation was about 1000 m. The importance of this quartz diagenesis to the structural history of one sedimentary trough and the reconstruction of the history of oil migration have been discussed by Philipp et al. (1963a,b). The sections that are close to the sand-shale interface, have much higher amounts of quartz overgrowths (Fig. 3-226). This increase in quartz content on approaching the sand-shale boundary is present in all sections of oil-filled sandstones studied by Fiichtbauer and can be explained only by assuming continuation of quartz precipitation after oil impregnation, possibly at a reduced rate. It seems that silica migrated into the connate water of the oil-saturated sandstones from adjacent shales and siltstones. Inasmuch as the existence of large volumes of connate water is improbable in this case, diffusion must have played a role. The latter process always should be considered together with fluid movements caused by compaction. That silica migration from the shales actually occurred is demonstrable. Chlorite content increases with increasing depth at the expense of kaolinite and with liberation of silica. A mineral balance is presented in Table 3-LVIII (Fuchtbauer, 1967a). Transformation of kaolinite to chlorite upon burial from 1000 to 3000 m gives rise to 5.5% increase in SiOz content, which can be partly explained by the change of kaolinite to chlorite (4.3-1.1 = 3.2%) (see Table 3-LVIII). Additional silica might have been available from kaolinization of feldspar and from pressure solution of quartz at clay mineral contacts. (c) In the bottom part of Fig. 3-226, the sandstone is not filled with oil at the present time. The quartz diagenesis is so similar to that of the oil-filled rock (b), however, that a former oil impregnation is assumed. Fuchtbauer (p. 358) stated that one may find surprisingly high porosities in very deeply buried coarse-grained quartz sandstones. Here, pressure solution is more pro-

El k

TABLE 3-LVII Relationships among grain size, rounding and porosity (%) (after Fuchtbauer, 1967a, table I, p. 356)

Coarse-grained Fine-grained

z

8in

Profile A (Fig. 3-225)

Profile B (Fig. 3-225)

Profile C (Fig. 3-225)

median

rounding1 porosity

median

rounding1 porosity

median

rounding1 porosity

0.42 0.12

3.35 2.4

0.23 0.12

2.08 2.10

0.22 0.08

2.53 2.17

22 27

“he “median rounding” from Fig. 3-225 is listed.

27 26

26 30

: z U

2 0 z

2!

$

U

d 0

5

$

3

0

z

K.H. WOLF AND G.V. CHILINGARIAN

350

nounced in the fine-grained than in the coarse-grained sandstones, which confirms Weyl’s (1959)postulation. F’iichtbauer found that carbonate cementation (calcitic or dolomitic sandstones) is not uncommon. The “minus-cement porosities” (Heald, 1956) have been determined by adding the volume of the cement to the pore volume. This gives the total porosity if no cement were present at the particular stage of burial. The “minus-cement porosities” generally range from 30 to 35%, which correlates with a depth of 600-1100 m according to Fig. 3-222,A. Those carbonate-cemented sandstones with 40% minuscement porosity must have been cemented soon after deposition and prior to distinct compaction. This is also confirmed by the good preservation of labile heavy minerals in such layers (Drong, 1965), as a result of an absence of intraformational corrosion and leaching (Pettijohn et al., 1972). Such cemented layers are most common near the top and bottom of sandstone bodies adjacent to shales or mudstones. The anions are retained in the sandstones by filtration at the sand-shale contact (Fothergill, 1955). The degree of carbonate cementation in the center of sandstones are generally related to grain size. The carbonate-cemented sandstones are most common in the fine-grained layers that are overlain by, or intercalated with, coarse-grained sandstones (Fiichtbauer, 1967a,p. 358). Other authigenic minerals were also mentioned by Fuchtbauer. Authigenic kaolinite is often confined to the purest sandstones (e.g., left part of Fig. 3-227),where its growth was not hindered or masked by detrital clay minerals, whereas early diagenetic dolomite or calcite cementation may have hindered kaolinite precipitation (right part of Fig. 3-227).The pore fluids may also have a controlling influence, e.g., oil impregnation may hinder kaolinite crystallization. As shown in Fig. 3-227,the kaolinite is poorly developed in the shales, is of primary detrital origin, and can be easily TABLE 3-LVIII Silica balance during the late diagenesis of Jurassic shales (after Fuchtbauer, 1967a,table 11, p. 357) Depth (m)

Quartz (%)

1000 14.5 3000 20.0 Si02 content, % increase + +5.5 -

Feldspar (%) (estimated)

Kaolinite (%) (ca. 45% Chlorite (%) (ca. Si02) 28% Si02)

5 5

63 53.6

0

-4.3[=(63-53.6) 0.451

17.5 21.4 X

+1.1[=(21.4--17.5) 0.283

X

DIAGENESIS OF SANDSTONES AND COMPACTION

DOGGER BETA KAOLINITE KAOLINITE - FIRECLAY FIRECLAY Q DIT0,IN DOLOMITIC SANDSTONE + DIT0,IN SANDSTONE NEXT TO SHALE 0

”k10

.

,

20

c

; ”. ’ L 0

30

P E R C E N T

,

40 50 C L A Y

,. .-.;.-. 60

70

351

11

1

80

c 2 O p

of the clay minerals in the fraction <0.02 mm) and the total clay content (ordinate) in sandstones (left portion of graph) and shales (right) of the “Dogger beta” Formation from Hankensbuttel and Luben-West, Germany. The remainder of clay is illite. (After Fiichtbauer, 1967a, fig. 6,p. 358; courtesy 7th World Pet. Congr.) Fig. 3-228. Relationship between the porosity and permeability of the “Dogger beta” Formation from Gross-Hamburg, Meckelfeld, Vorhop, and Wittingen-South, Germany. These measurements are taken from W. Tunn as are those of Figs. 3-229, 3-230A, and 3-230B. (After Fiichtbauer, 1967a, fig. 7, p. 359; courtesy 7th World Pet. Congr.)

distinguished from the well-formed authigenic kaolinite of the cleaner sandstone$. Similar to the authigenic quartz, the kaolinite in these sandstones is concentrated close to the sand-shale boundaries, whereas diagenetic chlorite appears to be more common in the shales associated.with greywackes.

352

K.H. WOLF AND G.V. CHILINGARIAN

Von Engelhardt (1960,p. 83) has offered and discussed a formula in his book on the relationship between porosity and permeability for sands with a restricted size range. Fiichtbauer (p. 359)stated that the shape of the “point cluster” in Fig. 3-228can be explained as follows. The vertical variation results from grain-size differences, and according to Fuchtbauer (p. 359)the “fact that decrease in porosity and permeability does follow the drawn theoretical curves indicates that the specific surface is not higher in the denser samples than in the more porous ones. This is possible only because the clay content of the sandstones of the ‘Dogger beta’ is generally low. Porosity decreases only through denser packing because of greater depth of burial as well as by carbonate, silicate, anhydrite or pyrite cementation. Both processes, however, do not appreciably change the specific surface.” The points above the curve are indicative of samples cemented by pyrite, whereas those below the lower curve represent argillaceous samples. The data presented in Fig. 3-229is different in that the samples are from a narrow depth range and are uncemented. Porosities, therefore, deviate slightly to low values at the same permeability. There is a decrease in porosity with increasing permeability*, resulting from increasing grain size and roundness, which enhances mechanical compaction. According to Fuchtbauer, “the clay content (less than 20 p in size) is responsible for the position of the fine-grained samples (median size is less than 0.12 mm) in Fig. 3-229.It increases with decreasing porosity from 5 to 30%, thus increasing the internal surface area and, therefore, decreasing the permeability. The open circles in the lower part of the diagram represent relatively coarse-grained sandstone with clay seams from the uppermost part of the sandstone body.” The dashed line corresponds to the median grain size of 0.12 mm and the highest median found (0.42 mm) is used for calculating the data for the upper curve. The break between samples with median values higher and lower than 0.12 mm evidently lies in the fact that the clay content is different. As shown on the left-hand side of Fig. 3-222,there is a decrease in porosity with increasing depth for three groups of sandstones having different medians. These arenites are rich in dolomite and calcite grains and have calcite and clay matrices. Porosity decreases with increasing calcite cement, but is independent of the dolomite content. The decrease in porosity with depth is much more distinct than in many other sandstones, and the grain-

* k = 93/[5( 1+)2S,ZI

(3-12)

where k is the permeability; Q, is the porosity; and SO is the specific surface area, which depends on grain-size distribution and shape of the grains. In Fig. 3-228, the two solidline curves, which envelope most of the points, were calculated and plotted for two constant SO values. Accurate formulae relating porosity, permeability, and specific surface area have been presented by Langnes et al. (1972).

DIAGENESIS OF SANDSTONES AND COMPACTION

353

Fig. 3-229. Relationship between the porosity and permeability of the “Bentheimer”

Sandstone, Scheerhorn, Germany. (After Fuchtbauer, 1967a, fig. 8, p. 360; courtesy 7th World Pet. Congr.)

size effect is different from that in the quartz arenites in that the coarsest ones are also the most porous. The carbonates (average content of 15%)in coarser calcareous sandstones occur mainly as detritial carbonate grains (Fig. 3-222,A),and in the form of recrystallized calcilutite matrix which reduces the porosity in the finer sandstones. Figure 3-222 also illustrates that the average porosity of oil-saturated sandstones is higher than that of oil-free sandstones. Figure 3-230shows a clear relationship between porosity and permeability. If this relationship were due to an increased cementation, the points should lie on the smooth curve of equal specific surface area or even beyond its lower left part. The downward-increasing deviation of the data points from the curve, on the other hand, indicates that the specific surface area increases with decreasing porosity as a result of the sediments becoming finer grained and more argillaceous. The curves in Fig. 3-231resemble the curves (Bausteinschichten Sandstone) on the left-hand side of Fig. 3-222.There is a

354

K.H. WOLF AND G.V. CHILINGARIAN

Fig. 3-23OA. Relationship between porosity and permeability of the “Bausteinschichten” Sandstone (= calcareous - calcite + dolomite -, clayey arenites of the Tertiary Molase). Solid-line curve is for constant specific surface area. (After Fiichtbauer, 1967a, fig. 9, p. 360; courtesy 7th World Pet. Congr.)

different reason, however, for the dependency on the grain size. In Fig. 3-222,A, the Bausteinschichten sediments show an increase in calcareous matrix content with decreasing grain size, so that the net result is a lowering of the porosity, whereas in the case of sandstones presented in Fig. 3-231 it is the particularly high clay content that lowers porosity. As shown in Fig. 3-232, the clay content is lower than 10%only in coarse-grained sandstones (Md larger than 0.5 mm). In sandstones with less than 30%clay, more than half of the clay is authigenic kaolinite (Fig. 3-232; see also Table 3-LIX). In studying compaction, one has to distinguish four types of clay occurrences described by Fiichtbauer, i.e., (a) layers, completely separated from the sand; (b) seams, relatively well separated from the sand; (c) uniformly distributed clays, authigenic (late diagenetic) in origin; (d) uniformly distributed, detrital matrix and early diagenetic clay minerals; and (e) combina.

DIAGENESIS OF SANDSTONES AND COMPACTION

1

I

I

000

1

500

UPPER ,CARBONIFEROUS

1

. ..

355

I

I

1

Fig. 3-230B. Relationship between porosity and permeability of the Upper Carboniferous sandstones from the Emsland region, Germany. (After Fuchtbauer, 1967a, fig. 10, p. 361; courtesy 7th World Pet. Congr.)

tions of the above. The influences of diagenetic processes, e.g., compaction and migration of fluids, will be different in each case, as discussed by Fuchtbauer. In cases where organic material is present, mainly H20 (with dissolved humic acids) and COz are released during the first stage of coalification, causing an acid environment in which feldspars are dissolved and replaced by

U

0 I

l-

a W

0

POROSITY, 7-

Fig. 3-231. Relationship among porosity, maximum depth of burial, and median grain size in water- and gas-saturated sandstones of the Upper Carboniferous sandstones in western Lower Saxony. (After Fuchtbauer, 1967a, fig. 11, p. 362; courtesy 7th World Pet. Congr.) Each one of the horizontal lines (a t o i) corresponds to one drilling well; they were obtained from curves showing the relationship between porosity and median grain size for each borehole. Maximum burial probably corresponds to the present depth. In sections a, b, and g, however, greater depth was formerly assumed for geologic reasons. In the lower right, the 0.5-mm curve is compared with the “Dogger beta” quartz sandstones from Fig. 3-222. .

TABLE 3-LIX Mean clay contents (<20 p ) in relation t o the median diameter (after Fuchtbauer, 1967a, table 111, p. 361) Sandstone types

Median diameter (mm) 0.1

Platform: Dogger beta Valendis Molasse : Tertiary Upper Carboniferous

0.13

0.16

0.2

Mean clay content (5%) 2.5 12.5

1.8 2.5

12.5

10 20.5

22

1.3 1.6 8 19

1.0 1.0 6.5

17

DIAGENESIS OF SANDSTONES AND COMPACTION

-z

-90

I-

; a

-80

w

Y

g -70 0

%

W

2 -60 W

357

UPPER CARBONIFEROUS KAOLlNlTE KAOLINITE-FIRECLAY FIRECLAY a-METAHALLOYSITE CHLORITE + MICA 6zPERCENT FELDSPAR IN THE SAND FRACTION '20 p

0

m

(L

2 -50

1 %

& -40 W

a J

t -30

-I

L1

5 -20 W z z tUi-10

c

I

I

Fig. 3-232.Relationship among clay (<20 p ) content, median diameter, porosity, and the relative amounts of kaolinite, mica, and chlorite in clay fraction of Upper Carboniferous sandstones. (After Fuchtbauer, 1967a, fig. 12, p. 364; courtesy of 7th World Petroleum Congress.) The numbers inside the diagram represent feldspar content in percent. The ordinate shows the peak heights of the 001 mica line, the 002 kaolinite line, and the 004 chlorite line in percent of the s u m of these peak heights. Quantitative ratios, therefore, cannot be read from the graph. All values are for one well from about 2,500 m depth, except for the points in the lower graph.

+

-2.4% 2.5-4.9'l.

DIAGENESIS OF‘ SANDSTONES AND COMPACTION

359

kaolinite. This occurs particularly in coarser sandstones that have little detrital clay and are more permeable to fluids prior to kaolinite authigenesis. The numbers in Fig. 3-232 indicate that the feldspar content is lowest in the sandstones containing more kaolinite. Some kaolinite, however, may have formed independently of the feldspar breakdown. According to Fuchtbauer (p. 363), the chlorite in shales is largely of secondary origin, whereas in sandstones it is mostly detrital. In the latter case, some chlorite may be the product of alteration of a detrital matrix. During the second stage of coalification, coal seams only release chemically ineffective methane. This causes an increase in pH, probably leading to a slightly alkaline environment. During this stage, mica (sericite) instead of kaolinite replaces the feldspar. In two publications, Fuchtbauer (1967a,b) described the origin and regional variation of the sediments discussed above. Some of the information is summarized below, because similar approaches are a prerequisite for regional analyses of compaction and movements of compaction fluids. In Fig. 3-233 three fluvial dispersal patterns are illustrated: A. transport direction into the sedimentary environment of K-feldspar-rich coarser sand with tourmaline, apatite and rutile heavy minerals; B. dispersal is characterized by albiteoligoclase finer sand with a garnet heavy mineral fraction; C. dispersal pattern characterized by plagioclase finer sands without garnet. The rivers passed into the brackish-lagoonal parts of a saline basin. The milieu of the northern basin was brackish t o marine through geologic time. In the south there are mainly sandstones, whereas in the north there are predominantly siltstones with numerous interbedded, unsorted sandstone lenses. The sandstones and siltstones contain muscovite and chlorite in the ratio of 6 : 1 to 1 0 : 1. The higher chlorite content in the north may be due to the nature of the depositional environment. The following cements were present: quartz, anhydrite, dolomite to ankerite, calcite, feldspar, halite, clay minerals (muscovite, chlorite, mixed-layer illite-montmorillonite, vermiculite, glauconite), analcime, and barite. The distributions of some of the cements are shown in Fig. 3-233. Distribution of feldspar, quartz and anhydrite cements in the Middle “Buntsandstein” (Lower Triassic) sandstone. Each symbol corresponds to the mean percentage (by rock volume) of secondary cement in one well or surface section. The sections were obtained from boreholes, except for the islands of Helgoland, A3, A8, and the section immediately north of the Main River. The means are based o n thin-section estimations using samples of 0.1-0.9 mm median diameter combined with samples >0.2 mm median diameter, in order to avoid grain size influences. Isopachs of the total Bunter except rock salt are taken from Trusheim (1963) and Sorgenfrei and Buch (1964). A, B, and C are the main dispersal systems (transportation directions). Analc. = analcime; Nu,K = areas of secondary albite and potassium feldspar, respectively; GI = glauconite; M L = mixed layer illite/montmorillonite; V = vermiculite; b = baritocelestite, barite and celestite; H = halite (see Fiichtbauer, 1967a). (After Fiichtbauer, 1967b, fig. 1 ; courtesy Sed. Geol.)

K.H. WOLF AND G.V. CHILINGARIAN

360

TABLE 3-LX Contact strength of sand grains cemented by different minerals (after Fuchtbauer, 1967a, table IV, p. 366) Cementing material

Number of samples Mean (3) Standard deviation (s) Student-t-test

Albite Quartz Anhydrite Open

7 19 13 6

1.20 1.59 2.15 2.26

50.075 50.145 *0.21 2 0.19

*0.11 *O.l 50.19 k0.34

Fig. 3-233.According to Fiichtbauer (p. 366), the following cement paragenesis is the most common: (1)analcime, vermiculite, mixed-layer clays, and illite; (2)feldspar and chlorite; (3)quartz; (4)calcite; (5) anhydrite; ( 6 ) dolomite; (7) barite; and (8) halite. The above sequence of precipitation of the cements was determined from studying the fabric and measurements of “contact strength”. Based on Taylor’s (1950) method, the numbers of tangential ( a ) , long ( b ) , concavo-convex (c), and sutured contacts ( d ) of sand grains were counted separately for areas with albite, quartz, and anhydrite cement. The values of contact strength* are higher with tighter grain contacts (Table 3-LX). Inasmuch as contacts become closer with increasing degree of diagenesis and compaction, the relative age of cementing minerals can be deduced by comparing contact strength of sandstones that contain different cements. A very conspicuous relationship exists between grain size and type of mineral cement: quartz is usually present in finer sandstones, whereas anhydrite cement usually occurs in coarser ones. The supersaturation for quartz in the adjacent fine-grained and coarse-grained sandstones was equal, but because the quartz grains acted as nuclei for precipitation and fine sandstones had a larger specific surface area (i.e., cm2/cm3 of bulk material present), more SiO &m was precipitated in the fine-grained sandstones than in the coarse-grained ones. Inasmuch as the porosity was equal in both fine-grained and coarse-grained sandstones, the porosity w a s reduced much more in the fine-grained sandstones. The already existing difference in permeability was further accentuated, so that the anhydrite solutions of later origin moved preferentially through the coarse-grained sandstones and cemented them. Thus, the grain size may control the differential cementation

____

* Contact strength = ( l a + 2b + 3c + 4d)/(a + b + c + d )

(3-13)

where a is the number of tangential contacts, b is the number of long contacts, c is the number of concavo-convex contacts, and d is the number of sutured contacts.

DIAGENESIS OF SANDSTONES AND COMPACTION

361

that, in turn, could control the rate and degree of compaction. The content of albite cement is higher in the fine-grained sandstones, because they contain more detrital plagioclase than the coarser sandstones and, thus, more nuclei are available for the precipitation of albite (Fig. 3-234). In fourteen out of eighteen occurrences, dolomite cement is present in fine-grained sandstones. Chronologically, NaCl appears as one of the last cements formed and occurs predominantly in the fine-grained sandstones, because the coarser sandstones have been cemented in the meantime by anhydrite. The contents of all other cements are independent of grain size. The distributions of the individual cements show clear regional differences. Vertical differences in cementation in individual profiles are of a more local character, but in general it seems that the deeper parts of the sandstones are more extensively cemented. The distribution of secondary feldspar is illustrated in Fig. 3-233,a. South of the Na-K line, only K-feldspar overgrowths are present; albite is more predominant in the north. It may be no coincidence that the Na-K line coincides with the boundary between the fhviatile basin in the north and the brackish-saline basin in the south. The 30

I

1

I

I

MEDIAN DIAMETER, rnrn

Fig. 3-234. Relation between feldspar content and median diameter of sandstones, based on X-ray estimations by Mrs. Goldschmidt; samples crushed to <35 p. Open circles = albite near A3 (see Fig. 3-233); open circles with dots inside = albite near A8 and east of Hannover; solid circles = potassium feldspar in both areas. (After Fuchtbauer, 1967b, fig. 2; courtesy Sed. Geol.) More or less thick rims of authigenic albite are bordering detrital grains of both albite and potassium feldspar north of the Na-K boundary line in Fig. 233a. Small rims of authigenic potassium feldspar are restricted t o the south of this boundary, which corresponds to the boundary between brackish-marine environment to the north and fluviatile environment to the south.

362

K.H. WOLF AND G.V. CHILINGARIAN

depositional milieu has apparently determined the secondary feldspar genesis, i.e., the earliest cement formed. In the terrestrial fluvial milieu, in which mica and K-feldspar decomposition through weathering occurs, K is enriched in the pore solutions as a result of weathering and the K is adsorbed on the fine-grained fluvial sediments. During compaction, K is remobilized and moves into the sandstones to be precipitated as K-feldspar. Under the conditions existing in the northern basin, where evaporation took place, the pore fluids became enriched in Na content (the concentration was even higher than that of K in the southern basin), leading t o extensive albite genesis. Analcime distribution, which is also shown in Fig. 3-233,b, is similar to that of albite. It is most abundant in the coarser-grained deposits and is confined to the weakly-evaporitic basin sections. Although quartz cement is quite widespread in its occurrence (Fig. 3-233,b), it is more enriched in the fluvial deposits; this becomes more obvious on examining Fig. 3-233,c. The incoming fresh water displaced the influence of the saline milieu. Also, the pore solutions became poorer in the sulfate content, with the resulting decrease in the amount of anhydrite cement. The precipitation of the silica was able to continue. The origin of the silica can be determined through thin-section examinations. In clay-rich sandstones, as well as in siltstones, where two adjacent quartz grains are separated by a clay film, stylolitic contacts are common and indicate dissolution of silica. In contrast t o quartz-quartz contacts, the quartz-clay contacts permitted removal of the silica in solution without interruption, so that pressure solution is more intense in clayrich sandstones (cf. papers in Spec. Publ., 7, S.E.P.M., 1959). The dissolved silica then moved into more pure sandstones t o be precipitated on the cleaner quartz surfaces. Various considerations by Fuchtbauer (p. 367) indicated that the main silicification began when the sandstone had a porosity of 20-30% and was at a depth of about 1OOOm. Chemical compaction becomes apparent below a depth of about 1000 m. Anhydrite is the second most common cement and shows a distribution complementary t o that of quartz (Fig. 3-233,c). It is most common in the central part of the northern depocenter and is absent in the fluvial basin. Fiichtbauer estimated that the anhydrite was precipitated when the sediments were at a depth of 1000 m below the surface. He attributed cementation to large volumes of sulfate in pore fluids, which moved during compaction from the fine-grained sediments (clays) into the sandstones. The anhydrite was precipitated when the NaCl content of the pore fluids reached a certain limit necessary t o exceed the solubility of CaSO,, as a result of the temperature increase during burial and the pressure decrease upon expulsion of the fluids from the fine-grained units. These units were under pressure above the normal hydrostatic pressure of the surrounding sandstones. There are two possible sources for the calcium sulfate: (1)enrichment through

DIAGENESIS OF SANDSTONES AND COMPACTION

363

evaporation in the lagoonal basins, and (2)compaction fluids from the Zechstein evaporite deposits. The original gypsum recrystallized to anhydrite at a critical depth of 500-1000 my with liberation of CaS04-saturated water which could have migrated upward into the sandstones. Baritocelestite, barite, and celestite ( b in Fig. 3-233,c) show a distribution similar t o that of anhydrite. As shown by Fuchtbauer, many of the diagenetic alterations are directly related to mechanical and chemical compaction. Some of the chemical constituents of the cements are of local origin, whereas others appear to have been derived from more remote sedimentary rocks undergoing compaction. Blanche (1973) observed that particularly in fluvial and sabkha facies, the clay matrix is composed predominantly of illite together with kaolinite, chlorite, and traces of montmorillonite. These minerals are derived from the decomposition of feldspars and micas. The permeability in these rocks decreases with decreasing grain size, whereas the porosity remains unchanged, sorting being its most important controlling factor. Following the reasoning by Fuchtbauer (1967,a,b) and Taylor (1950), Blanche expected the textural changes with increasing compaction as outlined by these two researchers. The maximum values of porosity and permeability, however, appear to be little affected by depth of burial down to 4000-4250 m. Below this depth a rapid decrease was observed, i.e., the porosity gradient was 1.3% for each 305 m of burial, which is in agreement with values published for other sandstone formations. Blanche also found that the depositional nature of the sands set physical limits on the reservoir potential: optimum values tended to occur in eolian dune facies, whereas the lowest reservoir parameters were present in the fluvial and sabkha facies. Powers (1967) treated the release mechanisms of fluids from clay-rich sedimentary units (see also Burst, 1969, for another example). Although the present chapter is confined to coarser sediments, the properties and behaviors of claystones, mudstones, shales, and bentonites (= pyroclastic deposits altered usually to ‘montmorillonite-rich accumulations) cannot be completely ignored. At least a brief treatment is in order here, because the stages at which certain amounts of fluids are released from clayey deposits determine how, when, and in what quantities these fluids are made available to the sandstones during compaction. The fluids from the finer sediments may or may not be related to compaction, they nevertheless influence diagenesis in general, including mechanical and chemical compaction, as well as cementation, decementation, and oil and ore-fluid migration. Together with studies such as done by Hitchon (1968) on the total volume of sandstones, shales, and limestones within certain sedimentary basins, the data on the release mechanisms of fluids from the clayey and other sediments and the information on the compaction history of the sedimentary pile will enable the

K.H. WOLF AND G.V. CHILINGARIAN

364

investigator to calculate the amounts of water, hydrocarbons, and, possibly, ore fluids that may be available per cubic meter of rock. At least the total volume of fluids could be estimated. Some of these studies have already proved of practical value in estimating the total amounts of petroleum and ores present in certain regions. Powers (1967)offered a compaction history (Fig. 3-235),which is based on laboratory and field data and the current knowledge of clay mineralogy. His conclusions can be summarized as follows: (1)When montmorillonite clays are buried to a depth of about 3000 ft, most of the water is expelled, except for the last few bound layers between the basal clay surfaces. This bound water may comprise nearly 50% of the volume of the clayey rock, and, apparently, cannot be squeezed out by further increase of burial pressure. (2)At a depth of about 3000 ft, the effective porosity and permeability are essentially zero for the mudstone. The clays and fixed water are "wrapped-around" the sand and silt particles. (3)A change from montmorillonite to illite begins at a depth of about 6000 f t and continues at an increasing rate to a depth of about 900010,000 ft, where no montmorillonite is left. This alteration mechanism is accompanied by desorption of the last layers of bound water from the clays and their release as "free water". (4) The fluids are released suddenly from the montmorillonite in the deep

WT11.mCH

rmJ

IO*I"O.ILLU(I.

DM

ILLIIC llU1.

111.0

"L..

YD I l 0 L I " l T l

Fig. 3-235. Compaction history of different clay minerals when deposited in marine environment and its probable relation to release of hydrocarbons from mud rocks. (After Powers, 1967, fig. 3, p. 1246; courtesy Am. Assoc. Pet. Geologists.)

DIAGENESIS OF SANDSTONES AND COMPACTION

365

subsurface, maybe during the clay-mineral transformation(s) and origin of mudstone f issility . (5)This sudden release of fluids would be absent from illite and kaolinite deposits because of the absence of a mineralogic change. Normal fluid movements by compaction would have occurred earlier. (6)During the montmorillonite-to-illite change, desorption of interlayer hydrocarbon layers and, possibly, trace elements takes place. (7)Below the “no-montmorillonite level” (Fig. 3-235),small amounts of water continue to be lost from the shale as the montmorillonite fraction of the mixed-layer clays collapses to form illite. This occurs at about 14,000 f t in the Gulf Coast area. Magara (1973)studied (1)relationship between differential shale compaction and depth, (2)relationships among porosity, permeability, density, and depth, (3)relationship between clay composition and depth, (4) relationship between the rate of sedimentation and compaction, (5) amount of fluid expelled during compaction, and (6)porosity distribution in shales as an indicator of permeability of associated sandstones and coarse limestones. According to Magara, “shale porosity distribution in incompletely compacted shale zones also may be affected by the permeability and the extent of adjacent sandstone or carbonate rock bodies. A sharp decrease of porosity in shales close to such rock bodies would suggest that relatively large volumes of fluids have been expelled from the shales into the adjacent sandstones or carbonates. If this expelled fluid volume is large, the possibility of hydrocarbon accumulation in such sandstone or carbonate rocks is considered to be favorable.” This has been confirmed by Magara’s studies of Mesozoic oil and gas pools in western Canada.

Compaction and stratigraphic correlation (Conybeare, 1967) Conybeare (1967)considered the results of compaction on stratigraphic correlation and consequent sedimentological-environmental interpretations. Inasmuch as differential compaction of a column of deposits in a sedimentary basin depends, among other variables, on the proportions of clay and sand, Conybeare theoretically considered (pp. 334-337) stratigraphic crosssections composed of different sand/mud ratios (Figs. 3-236 and 3-237). With increasing depth of burial, the total thickness of a section is reduced and the attitude of the sandstone beds is modified. The latter changes from an original attitude to one with a compactional slope, which may influence the fluid migration. In his model, Conybeare assumed that although the mud/sand ratios differ in sediment columns A and A’ (Fig. 3-236),which contain 1150 f t of solid sediments each, the rate of sedimentation was the same. The same assumption applies to cross-section A’--A3. For practical

K.H. WOLF AND G.V. CHILINGARIAN

366

0 I

lo00

II

!i ZOO0 1

aow

SAME LAYERS BURIED UNMR 2500' CLAY SOLIDS (UPPER 3500') Qf COMPACTIONAL SLOPE FROM 1 W 2

SHOWING CHANGE

4000

uoo CLAY :1000'SOLIDS

CLAY :500'SOLIDS

SAND :150'

SAND :650'

CLAY.lOO0'SOLIDS

SAND :1000'

Fig. 3-236.Schematic representation showing the effect of compaction o n sections A-A and A 2 - A 3 when they are initially buried under 3500 f t of clay muds. Sedimentation rates, in terms of total effective solids (100 f t of sand is effectively 100 f t of solids), are the same in columns A and A l , and in A 2 and A 3 . Columns A , A l , A 2 and A 3 have different sand/mud ratios and show variable degrees of compaction depending on the amount and depth of burial of mud layers. Changes in slope from depositional (1)to compactional (2), shown in hatchures between lines 1 and 2, are referred to as the "bellows effect". (After Conybeare, 1967, fig. 5, p. 338; courtesy Bull. Can. Pet. Geol.)

purposes the sandstone framework was considered constant, i.e., uncompactable, which is probably not true in nature (see Chilingarian et al., 1973).If the volume of sand is assumed to be constant, the degree of change of compactional slope will depend on the mud/sand ratio and depth of burial. Figure 3-237shows the thicknesses of sand and mud in columns A , B and B1 of a cross-section through a sedimentary basin. The sedimentation rate at B and B1 is twice that at A , and column B1 consists only of clay. As to the stratigraphic correlation, the sands at a depth interval of 3000-3250 f t in section B are equivalent to the sands at a depth interval of 1750-1900 f t in section A . The compactional slope between A and B at a depth range of 1750-3250 f t is 90 ft/mile or about 1degree. Conybeare calculated (see his table 2, p. 342) that an initial deposit of 837 f t of clayey mud contains 500 f t of solids, so that the amounts of clay solids in Fig. 3-237given as 1625, 2250 and 4550 ft for the sections A, B, and B1,respectively, were originally 2678, 3766,and 7533 f t prior to compaction. But as shown in Fig. 3-237, the actual thicknesses of sections A, B,and B1 are 3000,5250,and 5800 ft, respectively. Subtracting the sand thicknesses of 650 and 2300 f t in sections

367

DIAGENESIS OF SANDSTONES AND COMPACTION A

1000

2000

3000

B‘

B

F

4000 CLAY

. 1625’ SOLIDS

SAND : csd 5000

CLAY : 2250’ S 6000

SAND : 2300’

O

L

I

D

-

A

CLAY : 4550’ SOLIDS

Fig. 3-237. Schematic representation showing a section through part of a sedimentary basin where columns B and B’ each contain twice as much effective solids as column A , but have different ratios of sand to mud. The “bellows effect” illustrated in Fig. 3-236 results in expulsion of fluids from muds into sands; updip migration of these fluids is shown by arrows. (After Conybeare, 1967, fig. 6, p. 339; courtesy Bull. Can. Pet. Geol.)

A and B, the actual thicknesses of the clay deposits in sections A, B, and B1 are 2350, 2950 and 5800 ft, respectively. Subtracting 2300 from 2678, 2950 from 3766, and 5800 from 7533 results in 328, 816, and 1733, which represent the number of feet of water lost during compaction of the three sections. As Conybeare (1967, p. 337) stated: “Although the sedimentation rate in sections B and B1 is the same, section B1 has lost an additional 917 f t of water in comparison with that of section B, and B has lost 488 f t of water more than A ” (Fig. 3-237). Arrows in Fig. 3-237 indicate the directions of possible initial, updip fluid movements within the sand beds, depending on sedimentation and compaction rates. Conybeare called the squeezing of water from the clays into the sands and the consequent changes in the angles of compaction slopes the “bellows effect”. Updip fluid movements commonly occur toward a landmass, but need not necessarily be so. Penecontemporaneous or subsequent tectonism can alter or even reverse the compactional slope. Inasmuch as it has been shown by Chilingarian et al. (1973) that sands are just as compressible as clays, the above-described assumptions and calculations of Conybeare (1967) should be carefully reexamined. Future research

368

K.H. WOLF AND G.V. CHILINGARIAN

work hopefully will resolve this problem and determine under what conditions sands can be considered uncompactable in contrast to conditions when they are as compressible as muds. Baldwin (1971), as well as Conybeare (1967), pointed out that when muds and claystones intertonguing with sands are compacted, the differential compaction (based on the percentage of clay plus silt within the units) causes progressive distortion of the geometry of the initial depositional surfaces. Both of the above-mentioned researchers discussed “decompaction”, which is an analytical technique useful in projecting or extrapolating back the stratigraphy to an assumed earlier condition, i.e., restoration of the original depositional stratigraphy prior to compaction. Compaction is commonly expressed as a change in porosity (= ratio of pore volume to bulk volume) or as void ratio (= ratio of pore volume to volume of solid grains). “Grain proportion” (= ratio of volume of solid grains to bulk volume; Robertson, 1966, 1967) is the complement of porosity. Inasmuch as the volume of solid grains remains constant during compaction, grain proportion offers a simpler frame of reference than porosity. Baldwin proposed that decompaction can be accomplished by multiplying the present thickness of a compacted unit by a “decompaction number”, thus restoring the stratigraphic contacts to their precompaction position. The decompaction number, D, is expressed as: (3-14) where he =-earlier thickness, hp = present thickness, (pp = present porosity, earlier porosity, G, = present grain proportion, and Ge = earlier grain proportion. The grain proportion and porosity are expressed as decimal fractions in eq. 3-14. An approximate value of Gp can be obtained by measuring the dry bulk density of a sample in the laboratory and equating this to grain proportion in Fig. 3-238A, where the solid grain density is 2.66 g/cm.* Table 3-LXI presents the obtained data for North Sea deep-sea sediments (Hamilton, 1969a, tables 1and 2), and includes porosity and solid grain densities. At zero depth of burial (= precompaction stage), the value of Ge can be assumed to be 0.22 or 2276, which is equivalent to an initial porosity of 78% which corresponds to the weighted average of the data presented in Table 3-LXI. It is consistent with the data from Fig. 3-239, with a possible * 5 % @e =

* This is an average value used in soil mechanics (Hamilton, 1969b, p. 26) and is also the weighted average of data in Table 3-LXI.Most mud-forming minerals have an average solid grain density of 2.6-2.7 g/cm3 (Hedberg, 1936, p. 279).

369

DIAGENESIS O F SANDSTONES AND COMPACTION

0.5

1.0 1.1

BULK DIYUlY

a.0

a.s

1 DWTN Of DURIAL (fern0

I,./..)

VOID 14110

Fig. 3-238. Depth of burial graphs for clay and shale. A. Graph for converting bulk density t o grain proportion (and porosity). B. Depth of burial versus grain proportion (and porosity). For each curve, multiply depth of burial values by number shown for that curve. C. Graph for converting void ratio t o grain proportion (and porosity). (After Baldwin, 1971, fig. 1,p. 294; courtesy J . Sed. Petrol.) TABLE 3-LXI Porosity and solid grain density of some deep-sea sediments (after Hamilton, 1969a, tables 12) Sediment type

Continental Terrace Sand-silt-clay Clayey silt Silty clay

Number of samples

Solid-grain density Wee)

(shelf and slope) 17 2.71 40 2.713 17 2.69

Porosity average

(%I1

____-

standard error of the meana

67.5 75.0 76.0

1.66 0.87 0.74

Abyssal Plain (turbidites) Clayey silt 15 Silty clay4 35 Clay 2

2.61 2.55 2.67

78.6 85.5 85.8

1.53 0.49

Abyssal Hills (pelagic) 3 Clayey silt Silty clay 32 Clay 6

2.58 2.71 2.76

76.4 79.4 77.5

-

0.77 1.35

' Salt-free porosity; about 1%greater than without correction for dried salt (Hamilton, 1969b, pp. 25-27). Standard deviation = standard error X (number of Five samples (Hamilton, 196913, table D-1). Hamilton (1969b, table D-2).

*

370

K.H. WOLF AND G.V. CHILINGARIAN

variation in this initial porosity value, because the actual value depends on several variables. These include depositional process and rate and environment of accumulation, which determine the fabric, mineralogy, particle-size variation, organic content, and geochemistry of the deposit (Hamilton, 1969a,b; Meade, 1966). Figure 3-238Bpresents data for inferring the values of G, used in decompacting procedures to a condition of intermediate depth of burial. Where the value of G, has been obtained, the data can also provide the maximum amount of overburden and, as pointed out by Johnson (1950), the discrepancy between G, and the present overburden may indicate the amount of erosion at a surface of unconformity. The data in Fig. 3-238Bpertains to the approximate relation of grain proportion and porosity to depth of burial for clayey and muddy sediments, assuming that abnormal pore pressures were absent. This figure is flanked by conversion charts of bulk density (Fig. 3-238A)and void ratio (Fig. 3-238C).The average curve was obtained from the smoothed-out curve B in Fig. 3-239. 0

20

-x C

L

40

Z

2 c I

P z4

60

I

0

80

1oc DEPTH OF BURIAL ( f eat )

Fig. 3-239. Published porosity-depth data. See table 2 in Baldwin, 1971, p. 295. (After Baldwin, 1971, fig. 2, p. 295; courtesy J . Sed. Petrol.)

DIAGENESIS OF SANDSTONES AND COMPACTION

371

Using the above data, Baldwin (1971,pp. 296-299) proceeded to present examples of restoring compaction structures in interbedded shales (or argillites) and sandstones to their original precompaction geometric configuration. Although the above information is directly related to compaction of the clayey constituents in a stratigraphic unit, it has been presented here with the view that decompaction of the fine-grained units will also restore the interbedded sandstones to their primary, depositional, pre-compaction position, as demonstrated by one example, presented by Baldwin (see his fig. 4,p. 299).The stratigraphic and structural relationships between channel sandstone bodies, with their surrounding fine-grained accumulations, become much clearer after decompaction. As observed today (i.e., after compaction), the original outlines of the channel sandstones and the complex interbedding of the various lithologies can be visualized only with difficulty. Perrier and Quiblier (1974,p. 508) pointed out, however, that the decompaction number of Baldwin only relates to layers of infinitesimal thickness. Tmrnit (1968)has done extensive studies on stylolites (see pp. 133 and 139) and has discussed the various stages of pressure solution during the development of a subsiding sedimentary basin. For example, he presented a conceptual model (Fig. 3-240)on how the regional development of stylolites controls the movement of compaction fluids, an understanding of which is very important in the study of the origin of hydrocarbons and certain types of ore deposits. During the subsidence of the sediments in the center of the basin, stylolites are being formed in the carbonate rocks and, possibly, in quartzites. Carbonates require about 40 m and the quartzites about 1000 m

Fig. 3-240.Comparison between the unbedded or rarely-bedded massive limestones on highs (e.g., geanticlines) and the development of thinly bedded limestones in the basin (e.g., geosyncline). At the beginning, the dissolved material is able t o move vertically in the upward direction with the compaction fluids. As soon as cementation reduces the porosity, a strong vertical flow is eliminated. Consequently, the removal of any dissolved material in pore solutions occurs along the widespread pressure-solution surfaces in the direction of the structural highs. This material is then either precipitated as pore-filling cement or escapes into the ocean surface water. Solid arrow = flow direction of the compaction fluids; large open arrow = direction of basin subsidence. (After Trurnit, 1968, fig. 3,p. 382;courtesy Sed. Geol.)

372

K.H. WOLF AND G.V. CHILINGARIAN

of overburden for pressure solution to commence (see p. 378 in Trurnit, 1968). At the same time, the dissolved materials in the compaction fluids are carried upward into higher horizons where the pore space may become cemented and, consequently, the vertical fluid movements become increasingly impeded. The widespread occurrence of the pressure-solution surfaces, which ascend from the basin onto the geanticlinal ridges, then allows horizontal and diagonal fluid movements. The solutions, therefore, move towards the geanticlines and precipitation of cement takes place there, or the fluids are subaqueously expelled into the ocean water causing increased concentration of sea water in the chemical elements. Trurnit even suggested that microbrecciation of the carbonate sediments may be the result of expulsion of the compaction fluids at the flanks of the geanticlinal ridges. Migration of compaction fluids (Magara, 1968)

Magara (1968) studied migration of compaction fluids in marine Miocene mudstones and associated volcanic and pyroclastic rocks. His approach may serve as an example that can be applied in the future elsewhere. He found that the overlying and underlying mudstones were the source beds and that the water and hydrocarbons moved into the volcanic masses as a result of differential compaction. Magara (p. 2466) stated that it is important in petroleum exploration to determine the directions of movements and amounts of compaction fluids. One might add that this is also applicable to metalliferous ore exploration in cases where the deposits were formed by fluids of compaction and were controlled by stratigraphy. According to Magara (p. 2467),many investigators prepared lithofacies, biofacies, isopach, and subsurface structural maps to determine the migration and accumulation of hydrocarbons in a basin; however, no method that allows the determination of the direction of flow and the amount of the compaction fluids has been proposed. He, therefore, offered two methods to accomplish this, based on the horizontal and vertical porosity distribution, and porosity differences before and after compaction. A similar approach was used by Jobin (1962) in the study of Colorado-Plateau-type uranium deposits in sandstones. His porosity, permeability, and transmissivity data, however, were not related to compaction and, therefore, were not used to determine the migration directions and volumes of compaction fluids. Magara (1968)used the data available from the published literature (Fig. 3-241)on the empirically-determined relations between depth of burial and the degree of compaction of shales and mudstones. The curves show a marked decrease in porosity at shallow depths due to mechanical compaction. Under normal conditions, as compaction occurs and porosity decreases, fluids are expelled from the sediments. If the escape of the compaction

DIAGENESIS OF SANDSTONES AND COMPACTION

37 3

DEPTH, m

Fig. 3-241. Comparison of depth-porosity relations in several regions: Oklahoma (Athy, 1930); Venezuela (Hedberg, 1936); Gulf Coast '(Dickinson, 1951); Japan (Hosoi, 1963). (After Magara, 1968, fig. 5, p. 2473; courtesy Am. Assoc. Pet. Geologists.)

fluids is prevented or lowered, however, compaction may not be great and a high-porosity, high-fluid-pressure condition would result". Hence, more fluids would be expelled from rocks that can undergo compaction than those that do not. Magara stated (p. 2468) that, as a general rule, water is expelled from argillaceous beds into permeable carrier beds and then moves from localities of greater fluid expulsion to localities of smaller degrees of fluid expulsion. For this reason, it is possible to determine directions of compaction fluid movements from horizontal porosity distribution of the mudstone units above and below the carrier beds. Sonic and gamma-gamma logs (formation density logs) were used to determine porosity. Figure 3-242shows one example of porosity distribution of the mudstone overlying a reservoir composed of andesite agglomerate, and indicates differential regional compaction that resulted in controlling the directions of movement of the compaction fluids as shown by the arrows. In the present case, the arrows suggested the locations where oil or gas has accumulated. In other studies, maybe ore bodies could be located. Magara determined which beds acted as barriers and which ones underwent compaction. To calculate the volume of water expelled from the source rocks from a given time up to the Recent, one has to know the original thickness of the source beds. Magara calculated these thicknesses, after determining the normal porosity trend shown in Fig. 3-243.In Fig. 3-244,he reconstructed the formation thicknesses at the end

* For some data on high-pressure zones, see pp. 72-74.

i

Y,WI -WIT* lA

Y~NAMI-&UCHI I

.I

I

-

Fig. 3-242. Differential compaction map o f mudstone beds overlying agglomerates in the Nagaakaregion. (After Magara, 1968,fig. 8, p. 2477; courtesy Am. Assoc. Pet. Geologists.)

DIAGENESIS OF SANDSTONES AND COMPACTION

10

SHlUNJl SK-21 MUDSTONE ,POROSITY, 1

. 30

44

@

50

/

'

375

70% Uoourna Group

Hairune

I

Hahiyarn Famalia

-

Shiiya FOrmcllior

klKbju-m

FOrMtiW

N am lMi Farnatim

Shiinji Tuff

mx Fig. 3-243. Mudstone porosity values in Shiunji SK-21borehole. (After Magara, 1968, fig. 13, p. 2483; courtesy Am. Assoc. Pet. Geologists.)

376

K.H. WOLF AND G.V. CHILINGARIAN

of deposition of the Shiiya Formation, again by using his mathematical technique. The present porosity distribution after burial is also plotted in Fig. 3-244.The Teradomari mudstone just above a depth of 1340 m apparently has acted as a barrier to water movement, so that the water squeezed from the mudstones below this level would have moved downward and the water expelled above it moved upward to the surface and escaped into the sea. In Fig. 3-245,Magara illustrated how he used a simplified method of calculating the volume of compaction fluids. Assuming that the volume of solids remains more-or-lessconstant after compaction, the following formula can be used: V(1-

6 )= V‘(1-8)

(3-15)

where V = volume of sediments before burial, V‘ = volume of sediments after burial, 3 = average porosity before burial, fractional, and $’ = average porosity after burial, fractional. By determining first the after-burial, present-day vertical porosity distribution and then by assuming a normal porosity trend, the average porosity below the barrier before burial, 5,can be determined (Fig. 3-245).Also if the average porosity below the barrier after burial, T‘, and the after-burial, present-day volume of the sediment below the barrier, V‘, is known, the mudstone volume before burial, V, can be estimated (see formula 3-15): v = V”(1 --$)/(l -&)I (3-16)

It follows then, that the volume of compaction fluids, which moved downward, Vd, can be calculated:

v, = a v = (6- 6’) V’(.I -$)/(l - 5)

(3-17)

In applying his method over an entire region, Magara subdivided the area into many blocks and determined the total volumes of compaction fluids which moved downward within these blocks. Then, he was able to determine the directions of movements. From the present writers’ point of view, it seems that in the future such regional studies should take into consideration any data available on different compressibilities of various types of mudstones, sandstones, pyroclastics, etc. Fons and Holt (1966),for example, pointed out that montmorillonite shales (most probably a bentonite formed by the alteration of a tuff) are more resistant to compaction than illite and kaolinite. This, in turn, makes the montmorillonite also more resistant to internal fluid migration, which is related to the degree of water adsorption (see also Chilingar and Knight, 1960).As to the compactability of the different types of pyroclastic deposits, very little is known, but progress is being made at present (see Chapter 6 ) . The composition (i.e., vitric, lithic or crystal constituents), degree of zeolitization, clay neomorphism and other diage-

DIAGENESIS OF SANDSTONES AND COMPACTION SHIUNJI HUDSTONE

S1(-21

WROSllV,

0

377

70%

-Shiiya Farnoti0

-

Fa-0

Nmatanl M i

._--- F

Shiunji Tuff

mv. Fig. 3-244. Reconstruction of formation thicknesses at the end of the Shiiyan stage and the mudstone-porosity plot of Shiunji SK-21 borehole. (After Magara, 1968, fig. 14, p. 2484; courtesy Am. Assoc. Pet. Geologists.)

netic alterations, the rate of deposition, geologic time, and a host of other factors will determine the textural and structural characteristics of the tuffs, which in turn will influence their compactability. For an example on the calculation of the volumes of various lithologies (i.e., sandstones, shales, carbonates, and evaporites in this particular instance) within a sedimentary basin of western Canada and the total volume of fluids present, the reader may wish to consult the publication by Hitchon (1968). It should be remembered that Magara (1968) assumed a normal pressure gradient with depth. Some basins, or localities and units within particular basins, however, have abnormal pressures. Among others, Von Engelhardt (1960, p. 42) has pointed out that pore solutions are expelled. With a

378 MUDSTONE POROSITY

:.’

K.H. WOLF AND G.V. CHILINGARIAN

A S U M E D NORMAL POROSITY TREND

I IP W

D I-

t

W

w cn

a

i

RESERVOIR

Fig. 3-245. Diagram showing method of calculating volume of compaction c s r e n t . D ’ = thickness of mudstone after burial; V’ = Kolume of mudstone after burial; @ = average porosity below the barrier before burial; @’ = average porosity below the barrier after burial; and A@ = mean mudstone porosity decrease during compaction = (After Magara, 1968, fig. 15, p. 2485; courtesy Am. Assoc. Pet. Geologists.)

e@’.

decrease in pore space, the rate of porosity decrease with depth depends on pressure differences and the ability of clayey sediments to let duids pass (= permeability). Inasmuch as the claystone permeability is very small, the rate of compaction may not only depend on the rate of subsidence and increase

DEPTH, m

Fig. 3-246. Above-normal pressures in Tertiary sandstones in Louisiana. Solid circles = measured pressures; open circles = estimated pressures. (After Dickinson, 1953 ;courtesy Am. Assoc. Pet. Geologists.).

DIAGENESIS OF SANDSTONES AND COMPACTION

379

in the overburden pressure, but also on the rate of escape of the pore solutions. In clayey horizons that have not been completely compacted and where fluids are still moving upwards, the fluids in certain horizons must be under pressure which is higher than the normal hydrostatic pressure there. In the presence of interbedded porous sanstones the pressure will be particularly higher here. In the Tertiary clay-rich sediments of the Ventura Basin in California, at a depth of 2700 m the sandstone exhibits a pressure gradient of 0.23 atm/m (Watts, 1948). This pressure corresponds approximately to the weight of overlying sediments. In the Tertiary basin of the Gulf Coast of Louisiana abnormally high pressures have been also measured (Dickinson, 1953), as shown in Fig. 3-246. The lowest measured normal hydrostatic pressure was 0.107 atm/m, whereas the highest abnormal pressure was 0.21 atm/m. The highest pressures were present in sandstones within or beneath larger clay-rich units. A whole chapter was devoted by Rieke and Chilingarian (1974) t o the overpressured formations and their origin. Study on grain-contact types (Phipps, 1969) Phipps (1969) presented the results of a valuable study which showed: (a) the relative importance of grain contact types with burial; (b) the use of a “grain contact index”; (c) that compaction studies can assist in determining the time of cementation by chemical precipitation; (d) that differential compaction may be a function of differential chemical cementation from layer t o layer; (e) that the complex compaction history is a result of both cementation and decementation. A large proportion of Phipps’ data is given below. He studied an Eocene section of the Maracaibo Basin in Venezuela, comprised of a thick sequence of interbedded sandstones and shales. The former are quartzose, generally with more than 90% quartz grains. Chert may comprise up to 10% whereas the amount of interstitial clay is minor and seldom exceeds 5%by volume of sands. Micas, feldspar and pyrite are the most common accessories. According t o Taylor (1950), pore-space reduction and compaction can occur through: (1)solid flow of material under pressure involving grain fracture and distortion, and (2) solution and recrystallization-redeposition of material under pressure. Phipps stated that the sandstones studied by him display abundant evidence for mechanical processes of compaction, which include boehm lamellae, tension cracks in distorted quartz grains, displacement of twin lamellae in feldspars, cleavage along microfractures in feldspars, distortion and bending of micas between quartz grains, and “mushrooming” * of

* “Mushrooming” refers to the relationship between two grains where one became differentially dissolved and began to partly “wrap-around” the second grain.

380

K.H. WOLF AND G.V. CHILINGARIAN

quartz grains. Pressure solution and redeposition of silica cement or quartz overgrowths do not appear to be common and the reduction of porosity apparently took place mainly in response to mechanical processes of compaction. Although long contacts are common, secondary quartz overgrowths are rare. The relative ease with which the sandstones can be mechanically disaggregated indicates that cementation by secondary silica is not significant. The average total number of contacts per grain in a thin section is useful in assessing the degree of compaction (see Taylor, 1950, and section on textures in this chapter, pp. 156-163. This number is a function of grain shape, angularity, and packing, which are likely to be only of second-order importance as compared to sorting in sands composed of uniformly-sized particles. Inasmuch as a maximum value for the number of contactslgrain will be reached at some particular burial depth, the average number of grain contacts decreases rapidly in usefulness with increasing compaction, even for well-sorted sands. Below this depth, increased compaction will give rise to an increase in the number of concavo-convex and sutured contacts. Phipps found that the sorting has no influence on the type of grain contact. He has established an accurate technique for determining the degrees of compaction by employing the “grain contact index”. It gives a semi-quantitative value for the degree of compaction as based on type of grain contact. Phipps (p. 486) measured this index as follows: “The percentages of floating grains, and tangential, long, concavo-convex, and sutured contacts are multiplied by zero, one, two, three and four, respectively. These weighted figures, are then summed and divided by four to obtain the index. On this basis, maximum and minimum values are 100 and zero. Contacts per grain in thin section are derived by counting along regular traverses, the type of grain contact being noted at the same time. The number needed to get repeatable results is a function of the sandstone texture, and is best arrived at by making a few trial counts.” The findings of Phipps, who also examined in detail calcite- and sideritecemented sandstones can be summarized as follows: (1)Calcite-cemented sandstones occur as sporadic (3 inches to 1ft) bands within the main sandstone body. The sand grains are enclosed by the calcite crystals and are identical to those from non-cemented sandstones, but do not form the interlocking mosaic characteristic of the latter. In calcite-cemented sandstones, the boehm lamellae and other indications of mechanical compaction are very rare, whereas floating grains and tangential contacts are very common and the grain contact index and the average number of contacts per grain are very low (Table 3-LXII). In all other respects, these sandstones display the same grain size, sorting, and depositional characteristics as nonEemented sandstones immediately above and below them. Table 3-LXII

TABLE 3-LXII Variations of various properties and characteristics of Eocene sandstones from Maracaibo Basin, Venezuela, with depth of burial (after Phipps, 1969, table I, p. 487) Sample Depth No. (a)

Contacts Grain pergrain contact index

Floating Tangential grains contacts (%)

(40)

Long Concavocontacts convex (56) contacts

Sutured d m 2 contacts (mm) (5)

$0

3

Porosity Permeability (%) (md)

12755.5 12833.5 12834.0 12910.0 12976.5 13100.0 13230.0 13291.0

3.95 2.50 0.80

1.70 4.55 4.00 4.83 3.76

61 29 18 30 56 54 59 54

0

2.0 34.0 6.7 0

0 0 0

20.0 84.0 58.0 72.9 22.5 23.8 17.9 18.5

40.7 10.0 8.0 14.8 36.6 45.6 35.2 50.9

25.6 3.0 0 5.4 33.3 21.8 40.0 26.4

v1

Z

8

m

(%)

77 138 139 170 175 214 286 334

E b

13.7 1.0 0

0 7.6 8.8 6.9 5.2

0.259 0.230 0.180 0.137 0.202 0.230 0.185 0.286

1.25 1.29 1.36 1.46 1.20 1.20 1.24 1.20

12.8 11.4 16.5 3.9 15.3 14.1 10.6 13.9

121.0 25.0 122.0 <.01 112.0 55.0 8.1 368.0

Sandstones with carbonate cement; median diameter of framework fraction of sandstone (i.e., discounting clay fraction); So = Trask sorting coefficient of framework fraction.

2:

U m e l

0

z

*z8

U

2

2

w

00 U

K.H. WOLF AND G.V. CHILINGARIAN

382

shows that the uncemented arenites have undergone compaction, because the average numbers of contacts per grain and grain contact indices are twice as large as those of the cemented sandstones. No floating grains occur here and the percentages of concavo-convex and sutured contacts are high. On the contrary, in the cemented arenites at least 80%of the grains have tangential or floating contacts; concavo-convex and sutured contacts are rare. Porosities of calcitecemented sandstones are very low, but the percentage of carbonates plus percent porosity reaches a value of 35.1% (Table 3-LXIII). It can be deduced, therefore, that the cement was deposited prior to the occurrence of the main compaction. Inasmuch as the original depositional porosity was probably around 40%, there has been a porosity reduction of approximately 5%. This is in contrast to the porosity reduction of approximately 27% for uncemented sandstones. Cementation, therefore, was of early diagenetic origin and occurred after only a very slight compaction. (2) The large crystals of cement in siderite-cemented sandstones have been granulated along cleavage planes. These arenites are characterized by moderate to good porosity and permeability. Comparing Tables 3-LXII and 3-LXIII, it becomes clear that the siderite cement was also precipitated prior TABLE 3-LXIII

Variation of porosity, permeability, and carbonate content of sandstones with depth of burial1 (after Phipps, 1969, table 11, p. 489) Sample No.

Depth

Porosity

212 95 99 108 111 112 114 117 133 1382 13g2 166 169 170 172

12,692.5 12,767.5 12,770.5 12,776.5 12,786.5 12,787.0 12,795.5 12,797.5 12,817.5 12,833.5 12,834.0 12,858.0 12,868.5 12,910.0 12,936.0

14.6 3.5 7.3 3.0 2.2 1.9 2.1 2.7 7.7 11.4 16.5 8.5 7.5 3.9 4.7

1

Carbonate

ca.0.5 13.3 12.5 20.5 25.8 27.8 32.8 32.4 17.1 13.7 16.3 16.5 19.1 23.7 23.5

Porosity (%) Permeability (md) plus carbonate ca. 15.1 16.8 19.8 23.5 28.0 29.7 34.9 35.1 24.8 25.1 32.8 25.0 26.6 27.6 28.2

Percent interstitial clay is not accounted for in this table;

<0.1 0.1 <0.1 < 0..1 <0.1 <0.1 <0.1 <0.1 <0.1 25.0 122.0 <0.1 <0.1 <0.1 <0.1

siderite cement.

DIAGENESIS OF SANDSTONES AND COMPACTION

383

to the main compaction. In some cases, where the siderite cement has been almost entirely leached out, probably by subsurface water, leaving only small remnant granules of siderite in the interstitial spaces, compaction was able to take place. The minute granules, which are rounded cleavage rhombs, still display within the thin section optical continuity over areas up t o the size of the original larger crystal of siderite. Similar leaching is absent in the calcitecemented arenites. The greater porosity and permeability present in the siderite beds is apparently the result of leaching. Phipps (p. 488) reasoned that if the siderite cement was precipitated prior to compaction, then the siderite-cemented bands should have “frozen” the original textures and physical characteristics as the calcite-cemented ones did. Also, both types of sandstones should have similar very low values of porosity and permeability. Unexpectedly, however, the most obvious features of the cldexke arenites are their high porosity, high permeability, and granulated texture, which suggest a later, secondary origin for these features. As Phipps explained, if originally siderite-cemented sediment increased its porosity by the observed percentage, then this would necessitate a considerable expansion of the rock framework in order that porosity plus percent of carbonate cement could exceed the primary depositional porosity, i.e., porosity prior to leaching and granulation. As shown in Table 3-LXIII, however, this is not the case. The arenites with the maximum amount of unleached siderite have nearly the same percentage of combined pore space and cement as the calcitic sediments. Porosity and permeability must have been present to allow the passage of fluids that caused leaching. Phipps, therefore, proposed that calcite was present originally and that the secondary porosity was the result of sideritization, with a consequent volume decrease during replacement of the calcium by the iron ions. For similar possible changes, the reader may examine the data by Schmidt (1965). Inasmuch as complete sideritization of calcite leads to a volume decrease of about 20% (i.e., an increase in porosity of about 5% in an arenite with approximately 25% calcite cement), the process of sideritization invoked by Phipps does not account for the entire increase in porosity. The observed porosity is twice the amount that sideritization could account for. In addition, the replacement of calcium by iron was only partial. It seems, therefore, that leaching of siderite has been responsible for the additional increase in porosity. A stage must have been reached during leaching, however, when the cement was unable t o support any longer the overburden pressure, with a consequent occurrence of compaction. Prior to that, the arenite had its maximum post-burial porosity. Progressive leaching accompanying compaction reduced the porosity until the sediment was fully leached and compacted. Figure 3-247 illustrates the complete paragenetic sequence of the processes involved. Phipps assumed that sideritization of the calcite cement

384

K.H. WOLF AND G.V. CHILINGARIAN

zg

TIME (NoScale)

-

Fig. 3-247. Schematic illustration of the sideritization and leaching processes. (After Phipps, 1969, fig. 1; courtesy Geol. Mag.)

occurred at the present depth of burial of the rock, which at the same time is the maximum depth of burial. There is some evidence that this replacement occurred relatively recently. The sandstones with the greatest amount of remaining siderite cement are more porous than the texturally-identical, uncemented sandstones in the immediate surrounding localities. This suggests that the sediments are still compacting and that their higher porosity is of fairly recent origin, i.e., that leaching is probably occurring at the presentday depth of overburden. Sideritization, which had been an early post-burial event, was followed by subsequent leaching that allowed mechanical compaction processes to be initiated. The continued leaching, concomitant with increasing overburden pressure, maintained a porosity always slightly above that of the surrounding non-cemented sandstones, i.e., slightly higher porosity resulted from the time lag between leaching and compaction. The fact that sample 139 with siderite cement still retains a combined cement plus porosity percentage of 32.8 (Table 3-LXIII), which is of the same order of magnitude as the highest figures for the calcite-cemented sandstones, and that its “Grain Contact Index” is actually the lowest measured, is good evidence that compaction as a result of sideritization and leaching has hardly developed. That is, it has not yet suffered any significant compaction in addition to that which would be expected of a calcite-cemented sandstone. As Phipps pointed out (p. 492), “sideritization must have been a very recent event taking place at or near present depths of burial. In contrast, almost all the siderite cement has already been removed from sample 21” (Table

DIAGENESIS OF SANDSTONES AND COMPACTION

385

3-LXIII), “which only retains 0.5% of carbonate cement distributed as very small remnant granules. This sample has now compacted into a sandstone characterized by the same interlocking mosaic of sand grains common to all ihe non-cemented sandstones, and only differs from them in the retention of the siderite granules. Sample 21 represents the final end-product of the processes I have been describing, and it is clear that, if the siderite cement were not present, its history could not be distinguished from that of the non-cemented sandstones. Accordingly, it is also much more difficult to date the sideritization. The Grain Contact Index of sample 21 is 50 and contacts per grain are 4;both figures are in the same order as those for non-cemented sandstones at the same present-day depth of burial. . . Sample 21 illustrates an interesting point about the nature of the permeability of the sideritecemented sandstones. Despite the fact that sample 21 retains a relatively high porosity of 14.6%, its permeability is unmeasurable with normal methods in common with many of the non-cemented sandstones. This implies that some of the permeability of the siderite-cemented sandstone is a function of texture as well as porosity.” In regard to the mechanism of sideritization, Phipps believed that because the calcite-cemented sandstones have been impervious to interstitial fluids, the Fe-bearing solutions probably were not the cause of siderite formation and the nature of the fluids was changing from that of precipitating to that of dissolving the siderite. He proposed, therefore, that ionic diffusion is the most probable mechanism and that the immediate source for the iron was the pyrite present in all of the sandstones studied. Stylolites, which are very common, frequently contain highly ferruginous material, suggesting that the iron may have been mobilized during stylolite formation. This, in turn, also suggests that sideritization occurred at the present-day depth of burial. The publication by Cadigan (1971)might be mentioned here as an excellent example of very detailed qualitative and quantitative study of petrography and petrology of a formation on a regional scale. He investigated the Triassic Moenkopi Formation, which is one of the major uranium-bearing sandstones of the Colorado Plateau. Clastic grain types, heavy minerals, cement varieties, grain-size variation, sorting, etc. were studied on a regional basis and distribution maps prepared from the data. Similar investigations would be most useful in regional compaction studies. Alluvial fan deposits (Bull, 1972)

Bull (1972;see also Bull, 1964, 1973) undertook a study of land subsidence, compaction, and sediment-filled tension cracks in alluvial-fan deposits. During one of the civil engineering projects in California, thousands of cracks were found. They required investigation, because these fractures have

386

K.H. WOLF AND G.V. CHILINGARIAN

a direct bearing on the design and maintenance of structures, such as canals. According to Bull (1972),the land subsidence, which caused the fractures, is the result of four possible causes: (1)tectonic movements, (2)withdrawal of petroleum and gas, (3) artesian-head decline, and (4)compaction due to wetting of certain alluvial-fan deposits. The removal of subsurface water ( 3 above) has increased the applied stresses and, therefore, increased the normal compaction rate of the upper 1000-3000 f t of sediments. A subsidence rate of up to 1.8 ftfyear was the result. The differential settling occurred in some areas as a result of irrigation, which wetted the soil (4 above). Both overburden load and clay content affected the amount of compaction due to this wetting. An irrigation test was made on one of the fans. The surface of the fan first rose due to the expansion of the clay after water was first added, but as water penetrated to greater depths there was a net reduction in the volume of the deposit. The amount of volume shrinkage increased with depth due to increasing overburden pressure. The measurements on the amount of compaction due to wetting at different depths are shown in Fig. 3-248.There is a straight-line increase in degree of compaction to a depth of 100 ft. Thereafter, compaction decreases because of the higher natural moisture content below a depth of 125 ft, so that the addition of water results in a less pronounced change in the strength of the material and, consequently, less compaction. The effect of clay content on compaction due to wetting is shown in Fig. 3-249.For the samples tested it was concluded that: (a) Samples with less than 2% clay content compacted only slightly when wetted. (b) Samples with more than 30% clay content not only resisted compaction but also expanded by swelling. (c) Maximum compaction due to wetting occurred in water-laid sediments at a clay percentage of about 12%. (d) Resistance to compaction increases with increasing amounts of clay content (above 12%) when wetting takes place. The swelling of montmorillonite clays also reduces compaction. (e) The net compaction reached zero at about 30% clay content. Bull also studied the possibility of future near-surface subsidence in California, especially of large alluvial fans, which could damage the California aqueducts. Poland (1972) undertook similar studies. Figures 3-250 and 3-251 show a regional compilation of land subsidence from 1920 to 1972. More than 20 f t of subsidence due to artesian-head decline, sometimes at a rate of 1.8 ft/year, was measured. The compaction effects reached down as far as 3000 f t below the surface. In his most recent research results on the same area in California, Bull (1973)demonstrated that the estimated specific unit compaction (= compac-

DIAGENESIS OF SANDSTONES AND COMPACTION

387

tion during a time period, per unit thickness, per unit applied-stress increase) of the sedimentary pile in the northern subarea is four times that of the southern subarea. This indicates distinct variations in sediment compressibility. Bull demonstrated quantitatively that one-third of this compressibility difference is a result of lower prior total applied stress in the northern in contrast to the southern subarea. The remaining two-thirds of this difference are the consequence of variations in the water-expulsion rates in the clay-rich beds of different depositional environments. The northern area is characterized by flood-plain deposits composed of extensive sand beds with thin clay beds, resulting in a relatively faster dewatering under increasing effective overburden stress. The accumulations in the south, on the other hand, are

14

12

COMPACTION, 9.

0

CLAY CONTENT, 9.

Fig. 3-248. Effect of overburden load o n compaction due t o wetting. Inter-Agency Committee test plot B, after 42 months of operation, Arroyo Ciervo fan, California. (After Bull, 1972, fig. 3, p. 6; courtesy U.S. Geol. Surv.)

Fig. 3-249. Effect of clay content on compaction due to wetting under a simulated 50 ft overburden load. Squares = water-laid sediments; circles = mudflow deposits; triangles = deposits intermediate between mudflows and water-laid sediments; initial point on curve (ordinate) = Ottawa sand devoid of clay. (After Bull, 1972, fig. 2, p. 7 ; courtesy U.S. Geol. Surv.)

5311W P I

01

5

u

DIAGENESIS OF SANDSTONES AND COMPACTION

I ...I

I c . ~ surr., ”.,u sI*rso.&-l.Y

kmm

*4h,

389

1 c.ntr.i

Fig. 3-251. Land subsidence, 1926-197 2, LOBBanos-Kettleman City, California. (After Dr. J.F.Poland, personal communication.)

390

K.H. WOLF AND G.V. CHILINGARIAN

Experiments o f Pryor (1973) Pryor (1973) pointed out, as has been done by many others before him, that sandstones are the result of long and commonly complex histories of geologic evolution involving combination of: (1)sedimentary depositional mechanisms; (2)burial; (3) chemical and physical diagenesis (including compaction); and (4) structural deformation. In order to understand the complete history from the beginning to the end, a detailed knowledge of the primary (initial) grain accumulation characteristics is required. Little precise data has been collected in this area until recently by occasional investigations, in contrast to the more extensive information available on the secondary mechanical and chemical alterations, even though unresolved problems still exist in the latter field, too. Pryor (1973, p. 185) reported on some experiments with artificial sediments and stated that “the investigators jarred the sediments in their containers until no further compaction was observable and maximum packing was attained. Although natural sands do not receive this treatment, the winnowing processes of the beach and dune environments result not only in better sorting of the sands, but also result in a closer packing geometry. The ‘dumping’ processes of the fluvial environment yield a much looser, propped-open packing style than the beach-dune winnowing processes. Although no geometric data are available on the relative packing styles of sands from different depositional environments, packing styles may vary significantly and may be the cause of the textural and fabric differences shown in Table 3-LXIV. Von Engelhardt and Pitter (1951), among others, have shown that packing differences exert a strong control on variabilities in permeability and porosity. ” TABLE 3-LXIV Variation in carbonate content and composition, porosity, and grain density of Eocene sandstones from Maracaibo, Venezuela (after Phipps, 1969, table 111, p. 491) Sample Depth No. (ft)

Porosity

(%I

Carbonate Carbonate composition (%) (W wt) CaC03 MgC03 FeC03

Grain density2 (glcc)

117 139l 170

2.7 16.5 3.9

32.4 16.3 23.7

2.669 2.917 2.695

12,797.5 12,834.0 12,910.0

79.8 17.8 89.8

4.8 6.8 6.8

15.4 75.4 3.4

Siderite cement;2 grain density = total density of sandstone, including sand and cement, but not pore space, i.e., solids density not bulk density. Note markedly higher grain density of siderite-cemented sandstone.

DIAGENESIS O F SANDSTONES AND COMPACTION

100

400

boo

800

39 1

1000

Time (Seconds)

Fig. 3 - 2 5 2 . Effects of repacking by elutriation o n permeability and porosity of selected recent sands from different depositional environments. (After Pryor, 1973, fig. 25, p . 188; courtesy A m . Assoc. Pet. Geologists.)

To test the effect of depositional environment on packing, Pryor used duplicate samples collected from different sedimentary environments and allowed the sands t o be repacked through elutriation. Figure 3-252 demonstrates the change in permeability and porosity through time of elutriation as repacking takes place. In all cases, an initial rapid reduction in permeability during the first hundred seconds of the experiments was observed. This was followed by a levelling-off and stabilization of the rate of packing. Porosities also exhibit a marked reduction from the undisturbed to repacked grain arrangements. The data also illustrated that river-bar sediments change most extensively in packing in contrast t o beach and dune samples which change the least. Table 3-LXV summarizes some of the data collected by Pryor on river-bar, beach, and dune samples. He also concluded (p. 1 8 5 ) : “(1)Permeability increases with increasing grain size, and ( 2 ) porosity increases with increase in sorting; both relations agree with the experiments of Fraser (1935) and the determinations of Krunibein and Monk (1942).

TABLE 3-LXV

0 W

lu

Relations of permeability (k)and porosity (@)of sands to textural parameters and depositional environments (after Pryor, 1973, table 2, p. 188) River Bar

Grain size increases Sorting increases

Beach

Dune

Fraser (1935)

k

9

k

9

k

@

k

@

increase increase

increase increase

increase decrease

decrease increase

increase decrease .

decrease increase

increase decrease

decrease increase

x x

DIAGENESIS OF SANDSTONES AND COMPACTION

39 3

The relations of porosity versus mean size and permeability versus sorting separate the data into two different groups: (1)the river bars, where porosity increases with increasing sorting and permeability increases with increasing sorting (the reverse of Fraser’s, finding), and (2) beaches and dunes, where porosity increases with decreasing mean grain size and permeability increases with decreasing sorting (both in accord with the findings of Fraser)

...

9,

Pryor also described the permeability characteristics of bedding units of the beach, dune, and river-bar sediments; Fig. 3-253 depicts the latter. In river-bar sediments, there is a high variability in textural characteristics, permeability, and porosity, as well as many truncations and intersections as a result of cut-and-fill sedimentation. The permeability increases downward in each dipping laminae due to textural changes. The directions in which permeability increases are generally parallel to the length of the river point-bar deposits. Sand bodies of the river-bar and beach environments, the sediments of which exhibit well-organized and different patterns of permeability variations, are compared in Fig. 3-254. Pryor (pp. 188-189) stated that “permeabilities in river-bar sands are generally higher and more variable than in beaches and dunes and decrease systematically downstream and bankward. In beach sand bodies, permeability values have a small variability, but are

Fig. 3-253. Diagram of river-bar laminae packet with permeability characteristics. (After Pryor, 1973, fig. 15, p. 181; courtesy Am. Assoc. Pet. Geologists.)

394

K.H. WOLF AND G.V. CHILINGARIAN

RIVER BAR

HIGH VARIABILITY

BEACH

LOW VARlABlLlM

Fig. 3-254. Generalized relation of permeability trends in river bars and beaches. (After Pryor, 1973, fig. 26, p. 189; courtesy Am. Assoc. Pet. Geologists.)

relatively low on the beach faces, high on beach crests, and variable on the beach berms. Permeability-variation trends are parallel with the length of river-bar sand bodies and perpendicular to beach sand bodies. Dune sand bodies show no distinctive trend of permeability variation. There are greater variations of grain characteristics and permeability-porosity within bedding and lamination subunits than between them; this is especially pronounced in river-bar sand bodies. Packets of laminae and beds in beach-dune sand bodies are more uniform in character than those in river-bar deposits. Clay laminae and other low-permeability units commonly are present between bedding and laminae packets in river-bar sand bodies. In all deposits the boundary conditions between sediment packets are important factors in determining the effective reservoir characteristics of the sand bodies because sediment packets surrounded by other units of lower permeability will have effective permeabilities influenced by, and largely determined by, the lower permeabilities of the bounding units and hence will not demonstrate their ultimate through-flow capabilities. The ideal reactions between permeability-porosity and textural characteristics that have been shown by various authors for artificially packed particles are demonstrated only weakly by beach and dune sands and are not demonstrated by river-bar sands. The different styles of natural grain packing between river-bar sands and beach-dune sands are probably the cause of these deviations from the ideal models.”

DIAGENESIS OF SANDSTONES AND COMPACTION LATE DIAGENESIS-EPIGENESIS-BURIAL

395

METAMORPHISM

In the sequence late diagenesis-epigenesis-burial metamorphism*, the epigenesis (the term catagenesis is preferred by some researchers) and metamorphism probably are outside the realm of compaction of sedimentary rocks. Inasmuch as the various mechanical and chemical processes related to compaction grade into these stages, however, the writers decided to present a brief discussion on the burial processes and products beyond compaction. That there is no definite boundary between diagenetic compaction and epigenesis-metamorphism, can be easily seen by the progressive stages from orthoquartzite to a genuine metaquartzite, as well as by the transition from a clayey arenite through a chlorite-rich greywacke into a greywacke-schist. With these progressive changes in mineralogy, one would expect to find also changes from an uncompacted to compacted and, finally, t o recrystallized I

40,

TEMPER ATU R E ,a C

Fig. 3-255.Ranges of metamorphism. Conditions represented by points in the blank field on left are unlikely to be met in nature. A-B = range of measured geothermal gradients; C-D = range of geothermal conditions considered t o be common; E = melting range of granite in equilibrium with water vapor. 1 = diagenesis, weathering; 2 = dynamic metamorphism; 3 = thermal metamorphism; 4 = dynamothermal metamorphism; 5 = onset of melting; 6 = possible transient conditions in a geosyncline. (After Bayly, 1968,fig. 18-2; copyright @ 1968 Prentice-Hall, Englewood Cliffs, N.J.)

* It is important to note here that some investigators include epigenesis (or catagenesis, as preferred by some) in late diagenesis, whereas others consider it part of metamorphism. Some refer to the transitional stage between diagenesis and metamorphism as metagenesis (see Larsen and Chilingar, 1967). Some of the data presented here have been updated in the third edition of the book by Winkler (1967)published in 1974.

396

K.H. WOLF AND G.V. CHILINGARIAN

textures. Another example that illustrates the transition from diagenesis through epigenesis to metamorphism is supplied by the data on organic matter (“organic diagenesis-metamorphism”), which is discussed here also. The types of metamorphism as based on depth of burial or pressure and temperature are presented in Fig. 3-255.The processes related to diagenetic compaction and burial metamorphism belong to the shaded area depicting possible transient conditions in geosynclines. Two similar diagrams are given in Figs. 3-256and 3-257(Winkler, 1965)with supplemental data, especially important mineral associations. Figure 3-258presents data on the temperature variation in the Cenozoic sedimentary basins in the Gulf Coast of Texas and Louisiana. This data alone indicates that the late diagenesis and epigenesis may occur at temperatures near the boiling point of water and above, and that based on the temperature gradient there are no definite demarcations among diagenesis, epigenesis, burial metamorphism, and regional metamorphism. Fyfe (1973)presented the results of his study on dehydration reactions and stated that “the vapor pressures of hydrous materials indicate that thermal gradients in wet sediments may be influenced by endothermic dehydration reactions. Stress and salinity may influence the temperature of dehydration. The large solubility of stressed grains must lead to rapid elimination of porosity and to the condition that fluid pressures approach lithostatic pressures during sinking of wet sediments”. Examples, or case histories, of the various types of low-grade metamorphism are offered below, and the information on organic diagenesis is considered first. Landes (1967)called the incipient alterations of organic material “eometamorphism” in his investigations on the limits placed on the distribution of hydrocarbons by the intolerance of petroleum in the earth’s crust to TEMPERATURE, ‘C

Fig. 3-256.Schematic pressure-temperature diagram for different types of metamorphism. The ELT region below the lowest possible geothermal gradient of about lO*C/km is not realized in nature. The indicated depths corresponding to given pressures are maximum values and may be less than shown. (After Winkler, 1967, fig. 1 ; courtesy Springer, New York.)

L 66

NOIJ,L3VdpJO3 a N V SBNOJ,SaNVS 6 0 SIS3N3DVIa

m

398

W

(D

m

K.H. WOLF AND G.V. CHILINGARIAN

x

3:

3 G *z Fig. 3-258B. Temperatures in South Louisiana at a depth of 3048 m. Contour interval (CI) =, 5OC. Hot belt with temperatures between 100 and 113'C is near present shoreline. Rest of the area has temperature near 95 C. (After Jam et al., 1969, fig. 5, p. 2148; courtesy Am. Assoc. Pet. Geologists.)

U 0

. + Q

5

DIAGENESIS OF SANDSTONES AND COMPACTION

399

position between the other sedimentary rocks and petroleum in vulnerability to chemical and physical change brought about by increased heat and pressure . . . rank of the coal provides an excellent indication of the character of any hydrocarbons which may be stratigraphically nearby” (Hilt’s Law). Figure 3-259 represents an attempt to show the results of incipient metamorphism, as recorded by carbon ratios in coal and hydrocarbons, whereas Table 3-LXVI presents the relation between coal reflectance and hydrocarbon occurrence. The average relation between subsurface temperature and the density of hydrocarbons are demonstrated in Fig. 3-260. The amount of liquid hydrocarbons decreases whereas the percentage of gas and condensate increases with depth. Eometamorphism limits the occurrence of hydrocarbons in space both laterally in the basins bordered by mobile rims and at depth in deeper basins, due to increasing heat and pressure. According to Landes (p. 828), “it is estimated, based on past experience and therefore subject to modification in the future, that the commercial oil floor ranges from a depth

COAL

HYDROCARBONS Heavy oil and gos fields

?to Sub- Bituminous

OD0

000

Oil and gos fields

High volotile Bituminous c High volatile Bituminous 8

High volatile Bituminous A

Light oi! and gos fields Increasing percentage of gas fields Oil phase-out zone Gas oredominant

Gas only.

8o

I

Medium volotile Bitummous Low volotlle Bituminous

?---

decreoses downward

Semi- Anthracite Anthracite

00

Volume

Melo-Anthracile lo Graphite

lnodequote porosity

Graphite from enlropped gas

?

Fig. 3-259. Carbon ratios and occurrence of oil and gas. (After Landes, 1967, fig. 7, p. 836; courtesy Am. Assoc. Pet. Geologists.)

400

K.H. WOLF AND G.V. CHILINGARIAN

TABLE 3-LXVI Coal reflectance and hydrocarbon occurrence (after Ammosov and Syn-i, 1961, in Landes, 1967, fig. 8, p. 837) Stage

Vitrinite reflectance

~~

Coal type

Oil prospects

Lignitic Sub-bituminous Bituminous Bituminous Semi-anthracite Anthracite

Good Good Fair Gas only Nil Nil

~

1 5 8 11

60

74 83 92 125 170

17 22

of about 14,000 ft to 27,500 ft where the gradient is 1"/100ft. The commercial gas floor is an 'assay floor'; it is the level at which the volume of gas obtainable from reservoirs with decreased porosity does not yield a profit. It is concluded that commercial deposits of oil and gas do not extend to the basement rock in deeper basins." Staplin (1969) described progressive organic metamorphism, the next step beyond diagenesis as he pointed out, which affects particulate organic matter. He stated (pp. 56, 57): "The progressive changes include darkening of the organic matter, decrease in light transmissibility and electrical resistivity, increase in index of refraction, reflectance and lustre, and loss of fine structural detail. Increase in carbon ratio, spectrometric changes and other evidence of chemical transformation accompany the physical transformations .

..

200 LL 250 a 300

-1

!n -

w

0

a '

z

350 w 400+

I

2owo sIty and Gas Volume

Decrease Downward

1

1

Fig. 3-260. Earth temperatures and occurrence of oil and gas. (After Landes, 1967, fig. 9, p. 839; courtesy Am. Assoc. Pet. Geologists.)

DIAGENESIS OF SANDSTONES AND COMPACTION

401

The degree of organic metamorphism is determined by observing the organic debris, especially plant spores and non-woody cuticle and amorphous sapropelic debris. Of the organic materials in the rock, these seem to be the components most sensitive to heat, and their alteration in a clayey matrix progresses at roughly the same rate. Phytoplankton, less sensitive and more variable in its response, is less useful as a ‘thermometer’. The progressive severity of organic thermal metamorphism is subdivided into five degrees (thermal-alteration index, Table 3-LXVII), which are easy to determine for plotting on maps. As shown on this Table, sediments in areas with unaltered to moderately altered organic matter (thermal index 1-3) contain free and distillable hydrocarbons where the facies are suitable. In strongly to severely altered areas (thermal index 4 and 5) only dry gas is evident on the basis of cuttings, gas and hydrocarbon analyses, and experience gained through drilling shows that petroleum prospects in such areas, except for dry gas, are minimal.” Staplin studied (1)the distribution of gas in the Devonian carbonates of northeast British Columbia and equivalent areas in the Northwest Territories, and (2) the occurrence of oil and wet gas in the same carbonates in northwestern Alberta (Fig. 3-261). He found that the color and preservation of TABLE 3-LXVII Interpretative summary: type of organic matter, degree of alteration, and expected hydrocarbons (after Staplin, 1969,table 1, p. 57)

T y p e o f organic matter, matured facies (a) Sapropelic, amorphous (b) Plant cuticle, charcoal (c) Mixture of a and b Organic metamorphism (a) Blackening,:increase in index of refraction

(b) Pyrobitumen (c) Graphite, mineralization Thermal-alteration index (1)None (2) Slight (3) Moderate (4) Strong (5)Severe

Organic matter fresh, yellow brownish yellow brown black black, with additional evidence of rock metamorphism

Associated hydrocarbons wet gas and oil dry gas wet gas and oil wet or dry, depending on thermal alteration index

wet or dry wet or dry wet or dry dry gas dry gas to barren

K.H. WOLF AND G.V. CHILINGARIAN

402 WEST I CRETACEOUS MISSISSIPPIAN

z I WINTERBURN

EAST 2

3

4

5

6

7

8

D(I 11 1 f, 1

Fig. 3-261. Color of organic matter in eight wells along an e a s t w e s t section (northeastern British Columbia and northwestern Alberta). Color of organic matter: A = yellow; B = brown; C = dark brown and black. (After Staplin, 1969, fig. 2; courtesy Bull. Can. Pet. Geol.)

the organic matter, distribution of sedimentary hydrocarbons, presence of pyrobitumens, mineralization (e.g., galena, chalcopyrite), geothermal gradients and other variables all support metamorphism as the reason for the occurrence of dry gas in northeastern British Columbia. Staplin (p. 61)concluded: “Mineralization, high density and strong compaction, high-temperature gradients, clay organization, and deep depths of burial are evidence for thermal alteration, but the correlation between the appearance and properties of organic matter and the degree of metamorphism is the most rapid method of determining the effect of thermal activity in an area.” Baker and Claypool (1970) and Baker (1972)mentioned that incipient (low-grade) metamorphism leaves few obvious mineralogic or textural effects. Consequently, mildly-metamorphosed sedimentary rocks (other investigators may consider them epigenetic or late diagenetic in origin) are very difficult t o study. These authors also suggested that organic compounds are more susceptible to alteration and, therefore, may be more sensitive indicators of changes at temperatures and pressures existing between diagenesis and metamorphism than minerals. Thus, together with textural investigations, as discussed elsewhere in this chapter, it might be possible to develop a system of metamorphic facies classification for the realm of incipient metamorphism, based on the molecular and compositional character of organic substances. In Fig. 3-262,the hydrocarbon content is plotted against organic carbon content on a log-log graph paper. There is a positive correlation between these two variables for the unmetamorphosed samples, whereas no systematic relation is apparent for the metamorphosed samples. The hydrocarbon content of metamorphosed specimens is generally less than that of unmetamorphosed samples, and the difference is particularly evident for metamorphosed samples containing more than 0.5%organic carbon. There is

DIAGENESIS OF SANDSTONES AND COMPACTION ,J A x

403

UNMETAMORPHOSEO ”METAMORPHOSED” Martinsburg a equivalents Milligen a other Idaho samples

1.0 10.0 ORGANIC CARBON WT. %

100

Fig. 3-262. Relationship between hydrocarbon concentration and organic carbon content in metamorphosed and unmetamorphosed rocks. (After Baker and Claypool, 1970, fig. 1, p. 461; courtesy Am. Assoc. Pet. Geologists.)

a net loss of hydrocarbons during incipient metamorphism. Other data indicate that there is a relative increase of saturated compared to aromatic hydrocarbons as a result of incipient metamorphism. The relationship between saturated/aromatic and hydrocarbon/organic-carbon ratios in sedimentary rocks is presented in Fig. 3-263.Baker and Claypool (1970, p. 460) stated: “The unmetamorphosed samples fall into a linear trend with saturated/aromatic values less than 2 and a wide range of hydrocarbon/organiccarbon values. The metamorphosed rocks mostly fall in an area of low hydrocarbon/organic-carbon ratios and saturated/aromatic values greater than 2.” It is assumed that the metamorphosed samples were derived from material that originally would have fallen into the field defined by the unmetamorphosed organic matter. The metamorphic trend lines in Fig. 3-263illustrate the effects of incipient metamorphism on extractable hydrocarbons in sedimentary rocks: the absolute amount of hydrocarbons decreases, the ratio of hydrocarbons-to-organic carbon decreases, and the ratio of saturated-toaromatic hydrocarbon increases. The results of these investigations, therefore, indicate that the amount, composition, and molecular structure of the extractable hydrocarbons and the 3Ccontent of kerogenic carbon, according t o Baker and Claypool, can be used for the recognition and classification

HYDROCARBON/ORGANIC CARBON RATIO x

lo-’

Fig. 3-263. Effect of incipient metamorphism on extractable hydrocarbons in sedimentary rocks. Solid circles and dashed area = unmetamorphosed; x and stippled area = metamorphosed; triangles = Martinsburg Shale (Ordovician) and equivalents; x = Milligen Shale (Mississippian) and other Idaho samples (Paleozoic mudrocks); open circles = Triassic Lockatong Formation of New Jersey and Pennsylvania. (After Baker and Claypool, 1970, fig. 9 , p. 462;courtesy Am. Assoc. Pet. Geologists.)

+20 r +18. y+16

-

3 +14.

t +I2 +lo -

J

INCREASING METAMORPHISM MlLLlGEN

FM.

+8u ‘6’-

9

+2 -

+4

Fig. 3-264. 6C13 values of kerogenic carbon versus relative degree of metamorphism. (After Baker and Claypool, 1970, fig. 6, p. 465; courtesy Am. Assoc. Pet. Geologists.)

DIAGENESIS OF SANDSTONES AND COMPACTION

405

of incipient metamorphic facies. The general trend of the data obtained from the unmetamorphosed samples is controlled by the characteristics of the primary organic matter in the original sediments. Baker and Claypool (1970) concluded that (1) 613C values of the sediment extracts cannot be used as reliable indicators for either recognizing or classifying incipient metamorphism, and (2)the carbon isotopic composition of the total organic carbon of some metamorphosed clayey rocks may reflect the degree of metamorphism. Each of the groups of metamorphosed specimens used by Baker and Claypool (Fig. 3-264)show some qualitative progressive variation in 613C of kerogen carbon related t o the degree of metamorphic alteration, i.e., proximity to an igneous source or geographic location relative to the direction of increasing metamorphic grade, despite considerable spread of the data. In analyzing the data, the reader should be aware of various assumptions made by Baker and Claypool as well as of difficulties involved in the interpretation of the data. Kisch (1969) presented data on the relationship of some characteristic burial-metamorphic mineral assemblages to the rank of more or less co-equal coals (see also Frey and Niggli, 1971). He discussed: (1)alterations of clay minerals, and (2)appearance of zeolite-facies mineral assemblages. No details are presented here aside from the diagrams (Figs. 3-265-3-269),as they illustrate sufficiently well the principle of parallel mineral and coal modifica-

50

VOLATILE MATTER,d.o.f. Ye

Fig. 3-265. Derivation o f the vertical coal-rank scale from Karweil’s (1965, fig. 1; after Huck and Karweil, 1955) calculated volatile matter versus relative depth curve. (After Kisch, 1969, fig. 1, p. 410; courtesy Pergamon Press, Oxford.)

+=

MUSCOVITE AND ILLITE IN SANDSTONES

KAOLINITE

.

Fig. 3-266. Effect of late diagenesis on the distribution of clay minerals with depth in some deep boreholes. (After Kisch, 1969, fig. 2, p. 412; courtesy Pergamon Press, Oxford.)

?

cd

8

P

E z

*

h

i.

SIDERITE

t i

CULORITE IN CLAVSTONES

NVIIIV3NITIH3 ' A ' 3 CINV d?OM * H X

a

.i

j i

-

.a

CHLORITE AND MOTITE-CHLORITE IN SANDST.

!; - MUSCOVITE AND ILLITE \W CLAVJTONES

i

...

4

~HLORITE

.:

t

'

Aj

i

1

j

\

1

,

/ ...............

--

MONTMORILLONITE(l7~)

-

IIXED-LAVED ILLITE-MONTMORILLONlTE

..........

--

ILLITE(IOA)

M.-L. ILLITE-MONTMOR

Ll o o 0

90P

lb

0

UJ

DIAGENESIS O F SANDSTONES AND COMPACTION

407

tion with transition from the late diagenesis-epigenesis into low-grade metamorphism. Figure 3-265shows the vertical coal-rank scale (Karweil, 1956), as determined by calculations assuming a decrease of 2.2% volatile matter per 100 m increase in depth in the fat coal range and an average geothermal gradient of 40"C/km. In this figure, the volatile matter content is plotted versus the relative depth of burial and corresponding coal-rank parameter (based on volatile matter content). These relations hold for the Carboniferous Coal Measures of western Europe. Similar diagrams, however, should be prepared for areas with different coalification gradients. Figure 3-266demonstrates the diminution and, finally, disappearance during late diagenesis of the kaolinite and montmorillonite and the appearance of illite-muscovite and chlorite with increase in pressure and temperature. Figure 3-267presents schematically a correlation between the distribution of clay and carbonate minerals and the coal rank from Germany, Australia, and U.S.A. Kaolinite is PPRAMETERS

(vlrnlrr)

QUEENSLAND

MONSTE~LAND I M)PEWOLE. Wt5TWALIA

Fig. 3-267. Schematic distribution of clay and carbonate minerals with respect to the rank of associated coal (vitrite) in four late-Paleozoic areas. Coalification gradient is based on Karweil's (1956) curve (Fig. 3-265). Carbon contents for vitrites of stated volatilematter yield after Teichmiiller (1963, figs. 1, 2) and after Kotter (1960) (in brackets). Data schematized after ( I ) Kisch, 1968; (2) Scherp, 1963, Stadler, 1963, Esch, 1966; (3) Quinn and Glass, 1958; ( 4 ) Eckhardt, 1964, Teichmuller and Teichmiiller, 1966a, b. (After Kisch, 1969, fig. 3, p. 413; courtesy Pergamon Press, Oxford.)

K.H. WOLF AND G.V. CHILINGARIAN

408

I A M I C l r n

Fig. 3-268. Schematic distribution of the zeolites (analcime, heulandite-clinoptilolite, and laumontite) with respect to the rank of associated coal (vitrite). Coalification gradient as in Fig. 3-267. (After Kisch, 1969, fig. 7 , p. 419; courtesy Pergamon Press, Oxford.)

replaced by illite (muscovite), chlorite, and/or pyrophyllite with increasing rank of coals, but in different localities different changes in the clay mineralogy correspond to the varying coal ranks. As pointed out by Kisch, these discrepancies could be the result of differences in the primary mineralogy of the deposits and the availability of cations from labile detritus and circulating pore fluids. Figure 3-269indicates that during late diagenesis the kaolinite and montmorillonite disappear earlier in feldspathic sediments than in kaolinite-quartz-rich rocks. On the other hand, in the areas considered by Kisch (p. 415) “the burial-metamorphic disappearance of kaolinite in both these ‘rock families’ is associated with somewhat higher coal ranks (lean coal to meta-anthracite) than the coking coal to anthracite indicated by the correlation of mineral zones and coal types after Kossovskaya, Logvinenko and Shutov (1957)”(Fig. 3-269). Kisch (1969,p. 415) mentioned one case where the sedimentary rocks are

Fig. 3-269. Correlation of the distribution of some primary (stippled) and newly-formed (in black) silicate minerals in various terrigenous and volcanic sedimentary rock types during late diagenesis; compiled and schematized after Kossovskaya and Shutov (1961,1963).Broken bars indicate extension of some mineral zones (after Coombs, 1961). Correlation of mineral zones in terrigenous rocks with coal type after Kossovskaya et al. (1957).N o t drawn to scale. (After Kisch, 1969, fig. 5, p. 416; courtesy Pergamon Press, Oxford.)

410

K.H. WOLF AND G.V. CHILINGARIAN

rich in unstable detritus of igneous origin, which supplied alkalies and Mg and Fe ions. Here, the kaolinite disappears during the lean coal rank stages. Another example illustrates the absence of kaolinite and predominance of sericite-chlorite cements in sandstones and argillites associated with lean coals and anthracites (Fig. 3-269). The case of little or no availability of cations is represented by the kaolinite-coal tonsteins (Kisch, 1969, p. 415). He pointed out that in this case no noticeable replacement of kaolinite by illite and chlorite takes place as coal rank increases, i.e., the kaolinite persists up to anthracite rank, but its b-axis disorder progressively decreases up to a fat coal rank. Replacement by pyrophyllite may take place at the metaanthracite stage. Figure 3-268 illustrates that in the case of feldspathic-lithic and tuffaceous rocks, another group of mineral modifications occurs during burial metamorphism, i.e., the appearance of a succession of diagnostic zeolites and other Ca-aluminosilicates. With increasing depth of burial, these minerals tend to be progressively less hydrous and more dense. The coal-rank

Fig. 3-270. Precision of crystallite size measurements as a function of peak width. (After Griffin, 1967, fig. 1, p. 1008; courtesy J. Sed. Petrol.)

DIAGENESIS OF SANDSTONES AND COMPACTION

41 1

parameters and depth of burial have been presented along with the critical zeolite assemblages in Fig. 3-268. Griffin (1967) described an X-ray diffraction method to identify humic materials ranging in rank from lignite to meta-anthracite (or even graphite) as a result of diagenetic and metamorphic processes. Figures 3-270-3-273 present some of the results obtained by Griffin. Bishop (1972) studied sandstones containing pumpellyite and coal from relatively shallow depth. The inferred depths of burial suggest a pressure of not over 0.5-1 kbar and a temperature of not over 50--130°C. He presented

4

m

-01

Lignite

I roo

7

500

-02

400

Subbituminous A Subbturninous A Hlgh-Volotile 811 C

Hlgh-VolotlieBlt C High-Volotile Bd B Hlgh-VolOtlk Bit B

High-Volatile Bit A Medum-Volatk Bit

-03

Pr.

-04

200

-05 -06 -07 -08 -1 00 9

-20

-30 -40

-ao -60 -70

::1

-100

z I

' $

::

?ir

Low-Volatile Bit Sernionthrocite

- 200

Anthracite

-n o

- so0 a00

Mela-Anthrocite Grophite

Fig, 3-271.X-ray diffraction patterns of HCl, H F residues of the humic coal sequence (lignite to meta-anthracite). A pattern of standard graphite is presented for comparison. (After Griffin, 1967, fig. 2, p. 1008; courtesy J . Sed. Petrol.) Fig. 3-272.Variation in crystallite thickness and oxygen and hydrogen contents with coal rank. (After Griffin, 1967,fig. 3, p. 1010; courtesy J. Sed. Petrol.)

412

K.H. WOLF AND G.V. CHILINGARIAN

Fig. 3-273. C-axis dimensions of crystallites as a function of repeat distance (= d spacing) within the crystallites. (After Griffin, 1967, fig. 4, p. 1010; courtesy J . Sed. Petrol.)

an empirical straight-line curve of coal rank versus depth of burial for one particular area. From the rank number of the coal specimens associated with the pumpelIyite-containing sediments, Bishop was able t o determine depth of burial. In investigations such as those discussed above on organic matter, one has to consider the possible changes in composition that are the result of evolution and not due to secondary diagenetic or metamorphic modifications. Jackson (1973)observed that “humic” matter, isolated from bitumen extracted from suites of sedimentary rocks varying in lithology and depositional environment and ranging in age from Archean to Miocene, exhibited a systematic change in the molecular structure through geologic time. He attributed this change to major events in the evolutionary history of organisms. Jackson also found that despite diagenetic alterations, the primary differences between various humic deposits probably tend to persist under mild conditions long after burial. He did not comment on the possible effects of higher grades of diagenesis and metamorphism. A progressive change of the composition of organic matter with increasing burial depth and temperature was documented by Tissot et al. (1974). Murata et al. (1969,1972) discussed the isotopic changes of diagenetic

DIAGENESIS OF SANDSTONES AND COMPACTION

413

carbonates in terms of chemical processes that operate in deeply-buried marine sediments. This may also be a good approach in studying certain types of terrigenous sediments (shales, siltstones, sandstones, and conglomerates) if they contain carbonates in the form of fossils, cement, and/or concretions, as well as fracture fillings. If they can be shown to undergo changes with depth of burial, the information obtained is indicative also of the changes which occurred in the host rock. The alterations include changes in carbon isotopes, which are temperature-dependent, and in degree of dolomitization, sideritization, and pyritization, all of which can be influenced by composition of compaction fluids expelled from shales into coarser sediments. Murata et al. (1969)discussed certain isotope reactions between methane, generated by burial metamorphism of organic matter, and the carbonates that could account for the presence of abnormally heavy carbon in the carbonates. The effect of such reactions would be greatest in a deposit in which the ratio of organic shale to carbonate is large and in which the reacting fluids moved only along certain zones rather than by diffusion (Murata et al., 1969). Work on isotopes allowed them to conclude that the overall chronological succession of dolomite types, during the history of a sediment rich in organic material, seems to be: (1)lightcarbon dolomite, (2)heavy-carbon, low-iron dolomite, and (3) heavy-carbon, iron-rich dolomite. Types 1 and 2 belong to the long anaerobic phase of the sediment history, in the presence of organic matter in the sediments; whereas type 3 marks the first stage of the post-orogenic aerobic phase as a result of uplift of the formation. The latter allowed invasion of oxygen-rich surface waters, which released iron through oxidation of diagenetic pyrite. Thus, the minerals formed by fluids of compaction could be differentiated from those formed by post-tectonic fluids. Another example of the application of oxygen isotope study to problems of burial metamorphism is provided by Eslinger and Savin (1973). Anyone interested in the importance of compaction fluids in controlling chemical reactions, may wonder to what extent they have been influential in the origin of zeolites. Many publications on zeolite genesis mention the importance of such factors as (1)alkalinity of lake waters (e.g., Hay, 1966; Coombs and Whetten, 1967;and many others), (2) salinity of fluids (surface and subsurface solutions), (3) pH, (4) ionic composition, and ( 5 ) SiOz content. Little reference has been made to compaction fluids, however, most likely because they are very difficult to differentiate from other fluids that took part in the chemical reactions. Hay (1966)has shown that zeolites can form very early during diagenesis from volcanic glass reacting with alkaline lake water. Does that mean that should the latter be unavailable in the sedimentary environment, no zeolites can form? Or is there a possibility that at least later, during burial, saline compaction fluids, derived from another stratigraphic section, can move into

414

K.H. WOLF AND G.V. CHILINGARIAN

a volcanic glass-rich unit and give rise to zeolites? The answer to the latter question seems to be affirmative as demonstrated by late diagenetic and burial metamorphic zeolites. The composition of the fluids may be critical. Muffler and White (1969) reported on active metamorphism of Upper Cenozoic sediments of the Colorado River delta in the Salton Sea geothermal field in California, where a continuous transition from sediments through indurated sedimentary rocks to low-grade metamorphosed rocks of the greenschist facies occurs without the formation of zeolites. The authors suggested that the high aC02/aH20ratio prevented the formation of zeolites. For other examples of burial metamorphism, the reader may refer to the work of Crook (1963) and Coombs (1961). Packham and Crook (1960) presented a useful discussion on the problems related t o the transitional stages from halmyrolysis and diagenesis to metamorphism, and supplied diagenetic depth sequences. They pointed out one interesting difference between diagenesis and metamorphism: the original mineral composition of the rock is important in diagenesis, whereas the bulk composition is more significant in metamorphism. Coombs et al. (1959) assigned the zeolite facies to the lowest metamorphic stages, whereas Packham and Crook (1960) considered them as high-rank diagenetic facies because of the presence of recognizable essentially sedimentary textures. These facies owe their existence to the chemically-reactive volcanics. Thick, deeply-buried sediments lacking volcanics do not contain these authigenic minerals. There is a tendency of the zeolite facies to change with increasing age, as most of them are of Cenozoic or Mesozoic age and absent in Paleozoic and Precambrian rocks. Especially in eugeosynclinal sequences, pyroclastic sandstones may contain authigenic zeolites and other silicates that definitely show a progression of mineral facies as a function of pressure and temperature and, therefore, can serve as an index of depth of burial. It must be pointed out that certain types of zeolites can form under very low surface temperatures and pressures, i.e., those originating in saline lakes. For those readers who wish to gain information on the higher grades of progressive regional metamorphism, which influenced the chemical composition of a meta-arkose, Schwarcz’s (1966) paper can serve as a good starting point (see also Den Tex, 1965; Angel, 1965; Hietanen, 1967; Nelson, 1969; Ernst, 1971). It seems noteworthy t o mention Schwarcz’s suggestion that some of the reactions were isochemical in nature because (1)arkoses are composed of “inert” constituents (i.e., quartz, plagioclase, and alkali feldspar), which are stable over a wide range of temperatures and pressures, and (2) there is a lack of water of hydration to provide an aqueous fluid phase that, in turn, may have inhibited migration of chemical species and internal equilibrium during metamorphism. In contrast, the interbedded shales changed to pelitic schists which, according to Schwarcz, underwent many

DIAGENESIS OF SANDSTONES AND COMPACTION

415

stepwise reconstructive mineral transformations, with evolution of a waterrich fluid phase with increasing temperature. In this case, although compaction and compaction fluids had no obvious influence, the clayey material released fluids during metamorphism. Fluid release during diagenetic clay mineral transformations at much lower temperatures and pressures has been described previously in this chapter.

Zones of secondary alterations In terrigenous rocks of geosynclinal and cratonic (= platform) sedimentary sequences, Kossovskaya and Shutov (1958)established zones of secondary alterations on the basis of vertical change of various parameters, e.g., mineral composition and textures. Kisch (1969)used their results in his comparative investigations on coal rank. Three major zones (Fig. 3-269)were recognized, ranging vertically from the least to the most altered rocks, i.e., epigenetic, metagenetic, and metamorphic, with four subdivisions of the former two. The information applies to one locality with Mesozoic and Paleozoic rock sections, and the zoning reflects all stages of alteration in response to (1) slow basin subsidence, (2)effects of interstitial fluids, (3)increase in pressure and temperature, and (4)stresses in the geosynclinal area. The four zones are briefly described here.

Zone of unaltered clay cement. This zone is found on platforms and in the upper horizons of the marginal parts of geosynclines. It is up to 15002000 m in thickness and is composed of loose or poorly-indurated sandstones and water-saturated claystones, with original mineralogy and texturp being preserved. Gravitational compaction and consolidation increases with age in older rocks, as the bulk specific gravity increases from 1.4to 2.10 and porosity decreases from 40 to 20%. There is a variation in the mineralogical composition of the clay present as cement in the sandstones and clay in the claystones. This zone is characterized by the presence of carbonate, sulphate and, rarely, chloride waters. Some diagenetic reactions have taken place here, e.g., corrosion of ferromagnesians and feldspars. Zone of altered clay cement. This zone is present in the lowest parts of the platform regions, whereas in geosynclines the zone includes the fringe areas of folded strata. This zone is quite thick (up to 6000 m and more) and consists of shales and argillites having a porosity of 4-5% and less in the lower part of the zone. The recrystallization of the clayey matter and the neomorphism (newly-formed minerals) are characteristic of this zone. Montmorillonite disappears and kaolinite changes to illite (= hydromica), and the clay cement in the sandstone occurs as chlorite-illite aggregates. The excess

416

K.H. WOLF AND G.V. CHILINGARIAN

of Si02, resulting from the clay mineral transformation, forms quartz (and chalcedony?) micropatches. The paragenesis usually is colored Fe-chlorite, colorless chlorite, illite, and quartz. The mineralogy of the cement minerals depends on the original clay composition and the type of clastic grains, e.g., acid plagioclases are accompanied by illite cement and ferromagnesians predetermine the presence of chlorite. Quartzites with a primary kaolinite cement (chemically precipitated?) may remain unaltered except for some recrystallization of the kaolinite (for a similar change of illite crystallinity, see Ludwig, 1972), whereas in polymict sandstones the kaolinite changes to illite. The change of original clastic textures and fabrics is accompanied by quartz precipitation in the intergranular spaces. At a depth above 2000-. 2500 m, the subsurface fluids are highly mineralized and the solubilities of quartz and feldspars increase as a result of higher temperatures. Pressure solution is common here resulting in a supersaturation of the interstitial fluids with silica and consequent precipitation of quartz cement and formation of overgrowths on plagioclase. The transition from stage l to stage 2 can be recognized by the textural relationships of the neomorphic minerals, several examples of which were presented by Kossovskaya and Shutov (19 58). Zone of quartz-like structures and illite-chlorite cement. This zone is present in geosynclines only and did not develop in platform regions. This zone contains a complex of features determined by stress and depth of subsidence and is characterized by the presence of quartzites, schistose argillites, and clay-slates. Nearly complete change to quartzite and obliteration of primary clastic textures occur here. Grains show serrated contacts and composite blastesis. Silicification increases as a result of increasing temperature and pressure and even lithic fragments are recrystallized. Remnants of earlierformed cements are preserved as rims of chlorite and illite and are included in the blastic patches as separate entities. Where protective films or matrix is present to reduce the influence of interstitial fluids, however, the above alterations are absent. In this zone, the clay is altered to a chlorite plus illite assemblage and the detrital biotite undergoes a similar change. Zone of spine-like aggregates and muscovite-chlorite cement. This zone is developed in the central fields of folded regions and gradually changes to the upper zones of regional metamorphism. This zone is characterized by slates, phyllites and quartzites with complex spine-like structures, which are the result of interdigitation of chlorite and muscovite lamellae that penetrate the quartz and feldspar grains. The segregation of the two platy minerals is particularly evident in slates with large lepidoblasts of chlorite and muscovite set in a groundmass of scaly chlorite, muscovite and quartz. Illite is replacing

DIAGENESIS OF SANDSTONES AND COMPACTION

417

muscovite in various degrees depending on the stage of alteration, which is controlled by the loss of water and an increase in K content. The latter is derived from the decomposition of acid plagioclase and potash feldspars, as shown by corroded and muscovite-replaced orthoclase and microcline. The lower part of this zone consists of minerals that correspond t o the stable equilibrium of stress minerals at low-temperature regional metamorphism, i.e., chlorite, muscovite, quartz, albite, and epidote. Zones 1and 2 described above correspond to the so-called epigenetic stage and the dominant factors are vertical pressure and the influence of subsurface fluids. Zones 3 and 4 are stages of early metamorphism or “metagenesis” controlled by stress, vertical overburden pressure, and temperature. There is a gradation into regional metamorphism. (For a discussion on the various terminologies from diagenesis t o metamorphism, see Dunoyer de Segonzac, 1968, and Teodorovich, 1961.) Like the zones of regional metamorphism, zones l to 4 may cut across the stratigraphic boundaries. Sherwood and Huang (1969)studied highfy indurated carbonate rocks (limestones, dolomitic limestones, and dolomites) and igneous and metamorphic rocks (slates, schists, gneisses, diabases, and granites). They noted fundamental differences in the character of pores between the carbonate rocks and the igneous and metamorphic rocks (which need not necessarily be of general application in other areas): “Pores in the individual crystalline rocks are concentrated predominantly at two t o four different sizes, whereas those in the carbonate rocks generally are of one size. Also the mean pore diameters in the carbonate rocks are generally smaller; the pores are of a lesser size range than those in the igneous and metamorphic rocks, but the porosity of the carbonate rocks generally is higher and shows a greater variation. The low porosity and small pores measured in the dolomitic rocks from Virginia appear to indicate that most dolomitization preceded the strong compactive forces of Appalachian folding.’’ If studies such as these are performed on all major rock types in a sedimentary basin or in a region where mineralization in specific rock units has occurred, the results may suggest reasons for preferential or differential fluid movements and/or diffusion, as has been demonstrated in some cases already. As to diffusion processes per se, Elliott (1973)reviewed the theoretical data of the diffusion flow laws and included in his discussions pressure solution phenomena in sedimentary and metamorphic rocks. He found that microscopic observations could allow a clear distinction between deformation dominated by dislocation processes and deformation accompanied by mass transfer by diffusion. He concluded, however, that no microscopic criteria exist at the present time to recognize the path along which diffusion has occurred. As to detailed investigations of clay-mineral changes within sandstone

418

K.H. WOLF AND G.V. CHILINGARIAN Al,Si,O,,

A I z S ~ ~ O , o ~ O nH,O tI~~ idealized montmorillonite ( w i t h No, M Q )

(OH),

+si K AI,(Si,AI)

O,(OH),

ideolized i l l i t e

DEHYDRATION‘

Fig. 3-274. Possible dehydration among the aluminous clay minerals. Glauconite is chemically intermediate between the three molecules indicated. (After Bayly, 1968, fig. 13-10; courtesy Prentice-Hall, Englewood Cliffs, N.J.) Glauconite is of very early diagenetic origin, whereas other minerals may form as a product of late diagenesis and/or burial metamorphism within the sedimentary rock.

units during burial, much remains to be done, e.g., studies on the partition of chemical elements among coexisting phases. Figure 3-274 (Bayly, 1968) shows the possible diagenetic and epigenetic trends during burial in a sedimentary basin, excluding glauconite genesis of very early diagenetic origin; however, more complicated models with numerous intermediate stages can be set up, which is beyond the purpose of this chapter (e.g., Velde and Bystrom-Brusevitz, 1972). Another approach in which the length/thickness ratio of biotite is used, is discussed by Jones and Galway (1972)and Etheridge (1973).Although their approach was applied to higher-grade metamorphism, possibly one can use similar techniques to study minerals formed as a result of deep burial. Referring t o Kubler (1967,1968),Ludwig (1972)outlined a procedure by which the “crystallinity” of illite (i.e., the width of the 10 A illite peak at half height) is used in conjunction with index minerals t o determine the grade of regional metamorphism (Fig. 3-275).The X-ray pattern of illite clay minerals changes as a result of diagenesis and metamorphism. Thus, by using standards from illite-containing rocks of specific grade of secondary alterations, one can use the X-ray patterns to define the degree of secondary modifications. Four illite standards (i.e., Morris, Fithian, Marble Head, and OECD illites) were used after treatment with ethylene glycol (Fig. 3-275).The limit between anchimetamorphism sensu Kubler (= very low-grade metamorphism sensu Winkler) and epimetamorphism (= low-grade) is located at 4 mm along

DIAGENESIS OF SANDSTONES AND COMPACTION

419

-

“INCIPIENT TO WEAK METAMORPHISM:

r------------

’ A\

DIAGENESIS

I

OECD

I

‘\‘c SOmm

I

2

3

1

I

4

5

6

Fig. 3-275.Weaver’s “sharpness ratio” (= ratio of height of 10 8, peak to height of 10.5 8, peak; ordinate) versus Kubler’s “crystallinity” (i.e., width of the 10 illite peak at half height; abscissa: 0-20 mm, compared to quartz = 1-6 /A) for four illite standards in the grain fraction of 2-6.3 p. TS = Schwarzschiefer (black slate standard sample from the Staatliche Geologische Kommission, Berlin-Ost); TI3 = Tonschiefer (clay slate, source as for 2’s);1 = Morris illite; 2 = OECD illite; 3 = Fithian illite; 4 = Marble Head illite. Arrows indicate the oriented specimens treated with ethylene glycol. The limit between “anchimetamorphism” sensu Kubler (“very low stage of metamorphism” sensu Winkler) and “epimetamorphism” (“low stage”) is drawn at 4 mm. This corresponds to the value of 1.25, if the width at half height of the 4.26 8, quartz peak (3.2 mm) is taken as unit ( = 1.00). (Modified after Ludwig, 1972,fig. 1;courtesy Neues Jahrb. Geol. Paluontol.)

the abscissa. This corresponds to the value of 1.25, if the width at half height of the 4.26 8 quartz peak (3.2 mm) is taken as unit value (= 1.00). The above technique may also be useful in the study of progressive changes from syngenesis through diagenesis to burial and higher grades of metamorphism by using illite and other minerals, as demonstrated by Frey (1970), Frey and Niggli (1971), and Frey et al. (1973). The reader may wish t o refer to the publication by Velde and Bystrom-Bmsevitz (1972), who have experimentally studied the illite-montmorillonite evolution as a result of simulated burial metamorphism. The older studies of metamorphism of sedimentary piles have not as heavily leaned on the knowledge of the earlier diagenetic history of the sedimentary rocks as would appear to be desirable, so that, in a way, the metamor-

420

K.H. WOLF AND G.V. CHILINGARIAN

phism has been frequently considered somewhat “out-of-context”. Although the general composition of the pre-metamorphic rocks was always considered, more precise data is required in the future, and compaction must find its deserved place in these investigations. One example, provided by Huang and Wyllie (1973), is presented here. Although the petrologic and geochemical details fall outside the scope of this chapter, the information by Huang and Wyllie demonstrates, for example, the importance of interstitial water during metamorphism. The secondary processes and results thereof are different for sediments containing water as contrasted to those lacking fluids. Huang and Wyllie pointed out that plate tectonics, with its concept of consuming plate margins or sinking lithosphere slabs, can throw light on the styles of metamorphism occurring along the continental margins. Many of the intermediate and acid calc-alkaline magmas of the island arc and continental margins are derived from the slabs by progressive partial fusion of the siliceous sediments and oceanic crust of the slabs. They proposed experiments to solve some of the remaining problems: one should evaluate proposed natural processes by comparing the results of melting relationships of the theoretically inferred source rock materials (i.e., oceanic crust with sediments), on one hand, with those of the observed products of magmatic activity, on the other. The pressure range used would simulate that from the surface to great depth. The starting materials employed in their experiments were similar to many granitic plutonic rocks and may represent the lowmelting portion of some subducted oceanic sediments with a composition capable of metamorphism to muscovite-bearing quartzo-feldspathic rocks. Inasmuch as the compaction history of fine-grained and coarse-grained sediments determine the amount of fluids retained in a stratigraphic section, and because the sedimentary-volcanic pile will undergo burial and higher-grade metamorphism (related to plate consumption or not), the H20content will determine the paths of the secondary processes (compare figs. 1 and 2, in Huang and Wyllie, 1973). Consequently, the more complete and precise the available data is on the compaction history of sedimentary and volcanic deposits, combined with the information obtained from laboratory experiments, the more factual will be the theoretical reconstructions of metamorphism based on observed mineralogy, textures, and geochemistry. In their section of anatexis of metamorphic sediments in subduction zones, Huang and Wyllie pointed out that the following information must be available: (1)compositions and mineralogy of the rocks including water content and composition; (2) phase relationships of the materials present; and (3) temperature distribution in and around the subducted slab as a function of pressure, or depth. They proposed a model for magma generation with a zone of partial fusion, which gradually melts the lithosphere as the temperature increases. The first liquids originate from sediments or their

DIAGENESIS OF SANDSTONES AND COMPACTION

421

TEMPERATURE "C Fig. 3-276.A model for anatexis of subducted oceanic sediments in pressure (depth)temperature projection showing two episodes of melting. The curves b-u and b'-a' are the solidus curves from figs. 1 and 2 in the original paper. The dashed curves indicate published estimates of temperature distribution along the surface, and 1 km below the surface of a subducted lithosphere slab, according to Oxburgh and Turcotte (1970).The dotted curve indicates another estimate along the surface of a subducted lithosphere slab published by Toksoz et al. (1971). (After Huang and Wyllie, 1973, fig. 4; courte-7 Springer, Berlin.)

metamorphosed equivalents, whereas the oceanic crust will melt at greater depths. The paths for the surface of the slab and at a depth of 1 km are given in Fig. 3-276by the dashed line that bounds the stippled area, indicating when the sediments melt. Inasmuch as the solidus curve for material that contains excess water differs from the one for dry sediments (see figs. 1,2,and 3,in Huang and Wyllie, 1973), there are at least two solidus curves in Fig. 3-276 (i.e., a-b and ~ ' 4 ' ) . According to Garrels and Mackenzie (1971),the metamorphism of many ocean sediments could form metamorphic rocks composed dominantly of mica, quartz, and two feldspars, so that it is possible to use the phase diagrams for predicting the melting behavior of subducting sediments. The line a-b presents the conditions in the case of sediments with interstitial aqueous pore fluids, whereas line a'-b' is for sediments with all free water removed, possibly as a result of compaction. When subduction of sediments occurs with a lithosphere slab dipping at 45", then the temperature at the surface of the slab increases along the dotted line a-a' in Fig. 3-276.In the case where pore water was trapped in the sediments and was not removed by mechanical and/or chemical diagenesis, melting begins at point a, around 60

422

K.H. WOLF AND G.V. CHILINGARIAN

km below the surface. The first liquid would be saturated with water. Inasmuch as all of the pore fluids dissolve in a narrow temperature range, progressive fusion would result in a magma that is HzO-undersaturated (see fig. 3 in Huang and Wyllie). Although Huang and Wyllie have presented more information, the above discussion seems to be sufficient for the present purpose of pointing out that there are several gaps in the petrogenetic knowledge related to the boundaries (or transitional zones) between sediments and sedimentary rocks and sedimentary rocks and their metamorphic equivalents, which in turn change to igneous-like rock types. To close this gap in the genetic data, numerous models have to be established that include a complete paragenesis from early to late diagenesis, through low- and high-grade metamorphism, into the realm of origin of igneous rocks by remelting or fusion (= anatexis). A comprehensive study of compactional diagenesis would be required to determine the reasons for the amount of water retained in sedimentary-volcanic piles, as this quantity, in turn, will determine the fusion and mechanisms of igneous petrogenesis. STRUCTURES IN SEDIMENTARY ROCKS AS A RESULT OF COMPACTION

Enough data is available already on the origin of structures formed by the various mechanical and chemical compaction processes that could form the basis for a separate chapter. This is, however, beyond the scope of the present book, so that only a brief, alas very incomplete, list of references is given in tabular form without discussion, and the reader is referred to the corresponding references (see Table 3-LXVIII).

DIAGENESIS OF SANDSTONES AND COMPACTION

423

TABLE LXVIII Structures and related features formed or influenced by compaction Features formed by andfor influenced by compaction

References to pertinent literature

(1) Stylolites

Lerbekmo and Platt (1962);Golding and Conolly (1962); Trurnit (1968a,b, 1967); Schidlowski and Trurnit (1966); Park and Schot (1968) Franks (1969) Clifton (1965); Judson and Barks (1961) Gocht (1973); Fursich (1973) Rhoads (1970)

(2) Cone-in-cone (3) Polish on pebbles (4) Deformation of fossils by compaction (5) Reworking of sediments by organisms as a function of compaction; rheologic properties of sediments (6) Cleavage as a function of compaction (7) Concretions and their relation to compaction (8) Miscellaneous post-depositional structures influenced by compaction and rheology, e.g., load structures (9) “Quickstone” genesis (10) Mud lumps (11 ) Sandstone dikes (12) Faults formed by compaction (13) Sandstone geometry

(14) Surface topography due to compaction (15) Relationship between compaction and sedimentation, e.g., cyclicity (16) Dip changes due to compaction (17) Classification of sedimentary structures that consider compaction and rheology

Dunnet and Moore (1969); Alterman (1973); Powell (1972a,b) Oertel and Curtis (1972); Rukhin (1958); Fursich (1973) Dzulynski and Walton (1965); Swarbrick (1968); Elliston and Carey (1963); Wobber (1967) McNeill(l966); Boswell(l963) Dickey (1972) Steinitz (1970); Bull (1972); Shelton (1962); Conybeare and Crook (1968); Peterson (1968) Carver (1968); Phillips (1972); Powell (1972a,b); Bruce (1973) Pettijohn et a]. (1972); Baldwin (1971); Brown (1969); Rittenhouse (1961); Conybeare (1967); Bloom (1964); Kaye and Barghoorn (1964) Hrabar and Potter (1969); Dolly and Bush (1972); Halbouty (1972); Wobber (1967); Rittenhouse (1961); Brown (1969) Duff et al. (1967); Halbouty (1972) Rittenhouse (1972); Conybeare (1967); Borradaile (1973) Elliott (1965)

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