Pyrite-chlorite and pyrite-biotite relations in sandstones

Sedimentary Geology, 49 (1986) 177-192 177 Elsevier Science Publishers B.V., Amsterdam - Printed in The Netherlands

P Y R I T E - C H L O R I T E A N D P Y R I T E - B I O T I T E RELATIONS IN S A N D S T O N E S

SADOON MORAD

Department of Mineralogy and Petrology, Institute of Geology, Uppsala University, Box 555, S-75122 Uppsala (Sweden)

(Received November 21, 1985; revised and accepted June 3, 1986)

ABSTRACT

Morad, S., 1986. Pyrite-chlorite and pyrite-biotite relations in sandstones. Sediment. Geol., 49: 177-192.

Authigenic euhedral pyrite in feldspathic sandstones of the Visings~5 Group (Upper Proterozoic, southern Sweden) is observed to be closely associated with authigenic, fine-crystalline, interstitial chlorite

a n d / o r sand-sized chlorite or biotite. Microprobe analyses revealed that chlorites which are accompanied by relatively large amounts ( < - 3 vol.%) of pyrite are characterized by a lower Fe / (Fe + Mg) ratio than chlorites with lesser amounts ( < - 0.5 vol.%) of associated pyrite. The lower Fe / (Fe + Mg) ratio of the former chlorites might have resulted from subsequent substitution of Fe 2÷ by Mg 2÷ in an initially

Fe-rich chlorite; the released Fe 2÷ would then react with sulfide ions (in an anoxic-sulfidic environment) to form pyrite. Another alternative is that the chlorites with lower Fe / (Fe + Mg) ratio were crystallized from solutions characterized by ahigher activity of Mg 2÷ than chlorites with higher F e / ( F e + Mg) ratio. Iron from the biotite grains became available due to diagenetic illitization, kaolinization a n d / o r

chloritization.

INTRODUCTION

Pyrite, a well-known authigenic mineral from black and grey shales and organic- rich sediments, is formed by the reaction of sulfide ions with fine-grained iron minerals (e.g., Berner, 1970, 1981; Richard, 1974; Howarth, 1979). In a previous paper by the author (Morad, 1983), the importance of chlorite as a source of iron for pyrite authigenesis was presumed. Earlier, the reaction of fine-grained detrital chlorite with sulfide ions, which resulted in formation of authigenic pyrite in shales, was reported by Drever (1971), Siever and Kastner (1972) and Grossman et al. (1979). Siever and Kastner (1972) proposed the following reaction:

2(Fe3Mg2A1)(A1Si3)Ot0(OH)8 + 7CH20 + 4SO42 + 2Mg 2+ chlorite

= 2(Fe2Mg3A1)(A1Si3)O10(OH)8 + 2FeS 2 + 3H + + 7HCO 3 + 2H20 chlorite pyrite

However, none of the above authors have presented a mass balance relation for

0037-0738/86/$03.50 © 1986 Elsevier Science Publishers B.V.

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chlorite and pyrite arfd none demonstrated their detailed textural relationship, particularly by using scanning electron microscope.

The purpose of the present paper is to substantiate chemical and textural relationships between pyrite and interstitial chlorite, and between pyrite and sand- sized chlorite and biotite, in the feldspathic sandstones of the Visings~3 Group (Upper Proterozoic, southern Sweden). These sandstones belong to the upper part (about 60 m) of the middle unit (see Morad, 1983, 1986a) and were interpreted by Vidal (1976) as shallow marine deposits.

Thirty sandstone samples with variable amount of pyrite were selected for this study. They were examined petrographically and their modal compositions were determined by point counting (500 points per thin section). These sandstones are, in general, extremely rich in chlorite (Table 1). The microprobe analyses were performed on thin sections using an ARL-EMX instrument. An acceleration voltage of 15 kV, a sample current of 0.045-0.05/~A, and an electron beam diameter of 1-2 /~m were used during the analyses. The < 5 /~m fraction of sandstones was examined by X-ray diffraction analysis (see Morad, 1983). Chips of selected sample were investigated by a JEOL ® 840 JSM scanning electron microscope (SEM) equipped with an energy dispersive X-ray analyser (EDAX).

DIAGENETIC HISTORY AND THE TEXTURAL RELATIONSHIPS BETWEEN PYRITE AND

SILICATE MINERALS

Based on petrographic examinations, the paragenetic sequence of mineral diagenesis, from the earliest to latest stage, in the sandstones investigated is as follows:

(1) Alteration of detrital biotite and feldspars [by dissolution and/or replacement by small amounts of clay minerals (mostly illite)]. Detrital Fe-Ti oxides (ilmenite and titaniferous magnetite) were dissolved and/or replaced by hematite and titanium oxides.

(2) Precipitation of illite as thin coatings on the detrital grains (Fig. 1A). Presumably, this partially occurred simultaneously with stage 1 above.

(3) Precipitation of quartz, chlorite or both minerals simultaneously (Fig. 1A). In most cases, however, authigenic fine-crystalline chlorite solely filled the intergranular space (Fig. 1B). At this stage, chloritization of biotite and feldspar became a predominant alteration process.

(4) Precipitation of pyrite (perhaps partially simultaneously with some of the quartz of stage 3 above) in intergranular space (incompletely filled by chlorite of stage 3 above) and in the intercrystalline space between the chlorite flakes (e.g., Fig. 2C). Pyrite also precipitated in altered biotite grains (e.g., Fig. 4A). The Fe-Ti oxides were partly altered into pyrite and titanium oxides.

(5) Precipitation of coarse-crystalline low-Mg calcite. Thus, based on the sequence of diagenesis above, the pyrite appears to be formed

relatively late in diagenesis as compared to the well-known syndiagenetic-early

180

Fig. 1. (A) SEM micrograph showing illite ( I ) as coatings on detrital quartz (Q) and feldspar (F). The pores are filled by intimately intergrown authigenic chlorite (arrows) and quartz (q). (B) SEM micrograph showing a pore completely filled by pseudohexagonal flakes of chlorite.

181

Fig. 2. (A) Pyrite extensively filling the pore space in a feldspathic sandstone. Plane-light. (B) Details of the pyrite in (A) as shown by SEM. (C) SEM photograph of pore-filling pyrite (PY) surrounded by authigenic chlorite (Ch). (D) Poikilitic pyrite in which several grains of apparently corroded quartz and feldspar are embedded. Plane-light. (E) SEM photograph showing the development of a single euhedral pyrite (PY) within interstitial chlorite (Ch). Q is an authigenic quartz crystal.

182

Fig. 3. (A) SEM photograph of pyrite crystals (PY). Notice the pseudohexagonal chlorite crystals (ch) between the pyrite and a quartz particle (Q). (B) Fine-crystalline pyrite disseminated in interstitial chlorite. Plane-light.

diagenetic pyrites described by, for example, Berner (1970) and Howarth (1979). However, the above diagenetic stages have presumably been accomplished in the upper few tens of meters of sediments. According to A1Dahan and Morad (1986a), the sandstones studied have later been affected by a burial temperature of less than 100°C.

Pyrite occurs in relatively small amounts (up to about 37o) in the investigated sandstones (Table 1) and it has two major textural modes: (1) within interstitial fine-grained chlorite; and (2) within or attached to sand-sized biotite and chlorite.

Pyrite in textural mode 1 occurs as local patches and exhibits various textural forms. The most common form is a coarse-crystalline pyrite which fills the interstitial space extensively (Fig. 2A and B). In most samples chlorite lines the pore space, whereas pyrite fills Its central part. This texture is clearly demonstrated in the SEM photographs (Fig. 2C). Large euhedral and subhedral pyrite crystals (up to about 500/~m in size), in some cases, enclose several detrital particles, as well as chlorite coatings, thus generating a poikilotopic texture (Fig. 2D). Small, single cubic pyrite crystals (about 80-200 ~m in size) often appear to consist of several "micro-plates" arranged parallel to each other (Fig. 2E). The "micro-plates" vary in thickness and have smooth surfaces. Again, these crystals are also surrounded by chlorite. Pyrite was also found as numerous crystals (~ 30-60 /~m in size) within the interstitial chlorite (Fig. 3A and B).

The texture of the above chlorite suggests that it is of diagenetic origin (cf. Wilson and Pittman, 1977), because (1) it is characterized by well-developed pseudohexagonal crystals, and (2) it replaces and protrudes the detrital grains of feldspar and quartz. However, detrital chlorite matrix (intermixed with small

183

Fig. 4. (A) SEM photograph of a pyrite crystal (py) formed along the cleavage plane of a biotite (bi). (B) A pyrite crystal (PY) formed close to a chloritized biotite (ch). Notice the part of pyrite which conforms to a quartz particle (Q). Plane-light. (C) Fine pyrite crystals formed along cleavage planes of a chloritized biotite. Plane-light.

amounts of illitic clay as revealed from optical properties and microprobe analyses)

is present too. Such chlorites have a poorly developed crystal morphology (as

compared to the authigenic chlorite above). The detrital chlorite has usually been aggraded to chlorites similar to the authigenic ones (see Morad, 1984).

Mode 2 pyrite may occur as crystals of about 80-200 ~m in size which are embedded within, or attached to, biotite and chlorite grains (Fig. 4A and B). Frequently, the pyrite occurs as scattered small crystals (40 /~m in size) varying

greatly in number within the biotite and chlorite grains (Fig. 4C). The occurrence of authigenic Fe-rich minerals, particularly siderite, in altered biotite grains in a mode similar to the pyrites above has been reported by numerous authors (e.g., Bj~rlykke, 1984; Blackbourn, 1984; Huggett, 1984; Paveraro and Russell, 1984) in sandstone sequences they investigated. Such siderite was, however, not observed in the sandstones of this study.

184

MICROPROBE INVESTIGATION

Elucidation of chemical relationship between pyrite and interstitial chlorite was

carried out on six sandstone samples with good microscopic and X R D data. Three of these samples contain relatively abundant pyrite (Samples 3, 4 and 5; Table 1),

whereas the others (samples 14, 16 and 19) contain only small or trace amounts of pyrite. Microprobe analyses (Table 2) revealed that in samples containing minor

amounts of pyrite, chlorites have higher F e / ( F e + Mg) ratio (average about 0.76)

and A1 than chlorites in samples with higher pyrite content [average F e / ( F e + Mg)

is about 0.57]. However, all chlorites are iron-rich and have a total octahedral

occupancy lower than 6.00 of the ideal trioctahedral chlorite. This suggests the

possible presence of intermixing with dioctahedral chlorite. Furthermore, all of the

chlorites are characterized by higher octahedral than tetrahedral A1.

To elucidate the chemical relationship between pyrite and sand-sized biotite and

chlorite grains, microprobe analyses were performed on three each of such grains

which contain abundant pyrite as well as three with trace amounts or no pyrite

(Tables 3 and 4). The results showed that biotite grains with abundant pyrite contain less iron (average total Fe as FeO is 19.5 wt.%) than those with lower pyrite

content (average total Fe as FeO is 28.1 wt.%). Composi t ions of these biotites (Table

3) do not match the ideal biotite, because all have been subjected to diagenetic

replacement by clay minerals. Accordingly, microprobe analyses of such biotites

represent the average composi t ion of biotite and the submicroscopic clay-mineral

intergrowth irradiated by the electron beam (cf. Morad and A1Dahan, 1986a). The

most marked feature of these biotites is their low potassium content. The depletion

TABLE 2

Microprobe analyses of "non-pyritized" (14, 16 and 19) and "pyritized" (3, 5 and 4) chlorites. Formulae are calculated on the basis of O10(OH)s. Numbers correspond to sample numbers of Table 1

14 16 19 Average 3 5 4 Average

SiO2 29.1 30.3 29.4 29.6 32.9 30.3 30.9 31.4 AlzO 3 23.0 22.8 19.9 21.9 20.7 19.5 18.4 19.5 FeO * 27.6 25.7 24.3 25.9 17.3 20.2 21.2 19.6 MgO 8.2 7.2 9.4 8.3 14.6 15.8 13.6 14.6 K20 0.2 0.3 0.2 0.2 0.5 0.2 0.4 0.4

Total 88.1 86.3 83.2 85.9 86.0 86.0 84.5 85.5

Si 3.04 3.19 3.21 3.15 3.32 3.13 3.27 3.24 A1TM 0.96 0.81 0.79 0.85 0.68 0.87 0.73 0.76 AI vl 1.87 2.01 1.77 1.88 1.79 1.51 1.56 1.62 Fe z+ 2.41 2.26 2.22 2.30 1.46 1.74 1.88 1.69 Mg 1.27 1.13 1.53 1.31 2.20 2.43 2.14 2.26

Oct. occup. 5.55 5.40 5.52 5.49 5.45 5.68 5.58 5.57

* FeO represents total iron.

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TABLE 3

Microprobe analyses of "pyritized" (1, 2 and 3) and "non-pyritized" (4, 5 and 6) altered biotites.

Formulae are calculated on the basis of O10(OH)2


SiO 2 33.2 33.1 32.8 33.0 28.1 28.6 27.5 28.1

TiO 2 2.5 1.8 1.5 1.9 1.3 0.7 0.6 0.9

AI203 17.9 18.6 21.4 19.3 16.4 17.2 16.6 16.7

FeO * 20.1 21.2 17.2 19.5 26.9 27.3 30.2 28.1

MgO 6.4 9.6 7.4 7.8 10.8 9.7 9.0 9.8

K 2 0 5.2 6.0 6.7 6.0 2.2 1.9 2.3 2.1

Total 85.3 90.3 87.0 87.5 85.7 85.4 86.2 85.7

Si 2.78 2.65 2.67 2.70 2.43 2.47 2.41 2.44

A1TM 1.22 1.35 1.33 1.30 1.57 1.53 1.59 1.56

A1 vl 0.54 0.40 0.71 0.55 0.10 0.23 0.13 0.15

Ti 0.16 0.11 1.09 1.12 0.08 0.05 0.04 0.06

Fe 2÷ 1.41 1.42 1.17 1.33 1.95 1.98 2.22 2.05

Mg 0.80 1.14 0.90 0.95 1.39 1.25 1.18 1.27

Oct. occup. 2.91 3.07 2.87 2.95 3.52 3.51 3.57 3.53

K 0.56 0.61 0.69 0.62 0.24 0.21 0.26 0.24


is much greater in biotites with little pyrite, due to their advanced chloritization. This is supported by their high total octahedral occupancy (average of 3.53) compared to that of ideal biotite, with its 3.00 octahedral occupancy per O~0(OH)2. Chlorite can accommodate into its structure all, or most, of iron and magnesium

TABLE 4

Microprobe analyses of "non-pyritized" (1, 2 and 3) and "pyritized"

Formulae are calculated on the basis of O10(OH)8.

(4, 5 and 6) chloritized biotites.


SiO 2 29.4 30.1 29.8 29.8 28.5 30.9 30.2 29.9

A1203 17.2 18.1 18.1 17.8 21.2 19.7 17.6 19.5

FeO * 29.5 27.4 27.6 28.2 25.8 17.3 20.2 21.1

MgO 9.6 8.4 9.1 9.0 9.9 17.7 15.7 14.4

K 2 ° 0.9 1.0 0.7 0.9 0.6 1.1 0.8 0.8

Total 86.6 85.0 85.3 85.7 86.0 86.7 84.5 85.7

Si 3.21 3.30 3.25 3.25 3.05 3.15 3.21 3.14

A1TM 0.79 0.70 0.75 0.75 0.95' 0.85 0.79 0.86

AI vl 1.43 1.64 1.58 1.55 1.72 1.51 1.41 1.55

Fe 2.69 2.51 2.52 2.57 2.31 1.47 1.79 1.86

Mg 1.56 1.37 1.48 1.47 1.58 2.69 2.48 2.25

Oct. occup. 5.68 5.52 5.58 5.59 5.61 5.67 5.68 5.66


186

released from the parent biotite. As pointed out by Veblen and Ferry (1983), the chloritization of biotite may either occur by growth of brucite-like layers into the interlayer planes of biotite or by the replacement of TOT mica layers in the biotite by the brucite-like layer. On the other hand, the pyrite-bearing biotites were perhaps mainly altered to illite (and/or kaolinite), a process that is accompanied by ejection of octahedral Fe, Mg and Ti from the parent biotite. The presumed partial illitization of some of the biotites (analyses 1-3, Table 3) is indicated from their higher interlayer K and octahedral A1 contents, lower octahedral Fe and Mg, and lower total octahedral occupancies than the chloritized biotites (analyses 4-6, Table 3). Several authors (e.g., Gilkes et al., 1972; Goodman and Wilson, 1973; Gilkes and Suddhiprakarn, 1979) have related the expulsion of octahedral ions to the oxidation of Fe 2+ to Fe 3+. This oxidation process, the removal of octahedral iron, and the concomitant reorientation of (OH) groups in the biotite structure, result in an increase of the bonding strength of potassium (e.g., Barshad and Kishk, 1968; Farmer et al., 1971). A complete iilitization of biotite means that the mica structure changes from trioctahedral (biotite) to dioctahedral (illite). The two octahedral sites of iliite are predominantly occupied by A1 together with small amounts of mainly Fe and Mg. Illite in the rocks of the Visingsi5 Group usually contains considerably lower K than parent biotite (see A1Dahan and Morad, 1986a); thus part of K is liberated upon illitization of biotite. The diagenetic and low-grade metamorphic alteration of detrital biotite in sandstones has been discussed by White et al. (1985), Morad (1986b), and A1Dahan and Morad (1986a). From a study of alteration of detrital biotite in low-grade metamorphic clastic sedimentary rocks from the Sparagmite region of southern Norway (Morad, 1986), the formation of illite or chlorite is shown to be controlled by the chemistry of the pore solution. Illite was concluded to be preferably formed under conditions of higher H + activity whereas chlorite is formed under conditions of higher Mg 2+ activity.

Microprobe analyses of the sand-sized chlorite grains (Table 4) indicates that, as in the case for interstitial chlorite matrix (Table 2), those chlorite grains containing relatively high amounts of pyrite have a lower Fe/(Fe + Mg) ratio (average of 0.59) than chlorite grains with low or no pyrite content (average ratio of about 0.75). Microscopic examinations indicated that these sand-sized chlorite grains are altered biotite grains. They frequently have a higher birefringence than ordinary chlorites and sometimes contain preserved remnants of the parent biotite. The relatively high content of potassium is another indication of their derivation from biotite, because chlorites cannot accommodate potassium in their interlayer sites.

DISCUSSION AND CONCLUSIONS

Pyrite associated with interstitial chlorite

The lower Fe/(Fe + Mg) ratio which characterizes chlorites associated with high amounts of pyrite, might suggest that excess iron (required for the formation of

187

Envi ronlr~nt (A)

Oxic A

Anoxic- nonsulfidic A' (post-oxic)

Anoxic- A" sulfidic

S0#-,AI3+,Si4+,Fe3+Mg 2+

(Fel.69Mg2.26All.62)(AIo.76Si3.24)Olo(C~)8 + F~ ÷

(Fel.69Mg2.26AII.62)(AI0.76Si3.24)01o(0H)8

+ FeS 2

Environment (B)

SO~-,A]3+,Si4+,Fe3+,Mg2+

(Fe2.30Mgl.31All.88)(Alo.85Si3.15)Olo(OH)8

(Fel.69Mg2.26All.62)(Alo.76Si3.24)Olo(OH)8

+ FeS 2

Oxic

Anoxic- nonsulfidic (post-oxic)

Anoxic- sulfidic

Fig. 5. Two alternative models for explaining the mechanism by which iron became available for the authigenesis of pyrite and for predicting the origin of differences in Fe/(Fe+Mg) ratio in the fine-crystalline, interstitial chlorites (for further explanation, see text and Table 2).

(A) Chlorite characterized by a lower Fe/(Fe + Mg) ratio in samples containing relatively abundant pyrite has precipitated from a solution characterized by higher aMg2+ , but the same ave2+, compared to chlorites with the higher Fe/(Fe-Mg) ratio.

(B) The lower Fe/(Fe+ Mg) ratio in the "pyritized" chlorites was due to later substitution of Fe 2+ by Mg 2+ in an initially Fe-rich chlorite. Alternatives A and B would result in excess or liberated Fe z+, respectively, that could react with sulfide ions to form pyrite. The environments are after Berner (1981), A, A' and A" are envisaged as sediment beds of unknown thickness that could pass, upon burial, through the three types of environments.

pyrite) was available after precipitation of the chlorites. The availability of iron

might have either been achieved by (1) precipitation of the chlorites with a low

F e / ( F e + Mg) ratio from solutions characterized by a relatively higher aMg2+/aH+

ratio, but at a similar ave2+/aH, ratio, compared to chlorites with the high F e / ( F e + Mg) ratio (Fig. 5A), or (2) substitution of Fe 2+ by Mg z+ has occurred,

after the precipitation of an initially Fe-rich chlorite (Fig. 5B) (cf. Siever and

Kastner, 1972). Alternative (1) enables more Mg to be incorporated in the oc-

tahedral sites of the chlorite and would thus cause the lower F e / ( F e + Mg) ratio.

Assuming that all l iberated (or excess) iron reacted with sulfide ions to form pyrite, then the amounts of generated pyrite can be constrained by mass-balance calcula-

tions. The average compos i t ion of the "non-pyrit ized" chlorite is envisaged as a reference for initial compos i t ion and "pyrit ized" chlorites (Tables 1 and 2) as the

final products. The molar vo lume of chlorites was determined by the regression curve of Parry and D o w n e y (1982). The standard molar vo lume of pyrite is 23.94

188

cm 3 mole-a (Helgeson et al., 1978). Based on these data, the mass-balance equations (using the average chemical formulae of "non-pyrit ized" and "pyri t ized" chlorites) and the amounts of pyrite which could be roughly estimated from the percentages of chlorite in the sandstones, are as follows:

Sample no. 3 (Tables 1 and 2)

(Fe2.3Mgl.31A11.8s)(Alo.85Si3.15)O,o(OH)8 + 1.68S 2 + 0.17Si 4+ "non-pyritized chlorite"

+ 0.89Mg 2+ = (Fel.46Mg2.2omll.79)(mlo.68Si3.32)Olo (OH)8 "pyritized chlorite"

+ 0.84FeS z + 0.26A13+ pyrite

(pyrite = 2.5 vol.%)


"non-pyrit ized chlor i te"+ 1.12S 2- + 1.12Mg 2+

= (Fel.v4Mg2.a3All.51)(Alo.87Si 3.13 )O10 (OH)8 + 0.56FeS 2 "pyrit ized chlorite" pyrite

+ 0.35A13 + + 0.02Si 4 +



"non-pyrit ized chlor i te"+ 0 . 8 4 S 2 - + 0.83Mg 2+ + 0.12Si 4+

= (Fex.88Mgz.14All.56)(AI0.73Sia.27)O10(OH)8 + 0.42FeS 2 + 0.44A13+ "pyrit ized chlorite" pyrite


These calculations indicate that, on average, about 75% of Fe 2+ needed for pyrite authigenesis could have been accounted for by the difference in F e / ( F e + Mg) ratio between "pyri t ized" and "non-pyrit ized" chlorites. Other sources of iron may be detrital micas and Fe and Fe -T i oxide grains (see Morad and A1Dahan, 1986b).

The textural relationship between pyrite and interstitial chlorite (Fig. 2D and E) indicate that authigenic chlorite has been formed at an earlier diagenetic stage than associated pyrite. This indicates that sulfide ions became available after the forma-

189

tion of chlorite. Sulfide ions may have been formed by reduction of pore water sulfate, which escaped reduction in the chlorite zone because of the low concentra- tions and /o r poor quality of the metabolizable organic matter. Alternatively, solutions containing sulfide ions and magnesium, e.g., to substitute for Fe 2÷ in chlorite were introduced into the sandstones only after the formation of chlorite. The source of such ions is, however, enigmatic.

In summary, it is suggested that chlorite is formed in a sediment layer (layer A, Fig. 5) under reducing conditions (post-oxic environment; Froelich et al., 1979; Berner, 1981) where soluble iron became available (cf. Lynn and Bonatti, 1965; Li et al., 1969) probably by reduction of Fe 3÷ in layer A upon burial. Following further burial of layer A to a position of layer A", pyrite was formed due to the presence of sulfide ions in the anoxic-sulfidic conditions at this depth (cf. Berner, 1981). High Mg 2÷ activity must have been prevailed in order that Mg 2+ could replace Fe 2÷ in initially Fe-rich chlorites (Fig. 5B), or, more likely, to precipitate chlorites enriched in Mg (Fig. 4A). In the latter case, excess iron could have been available to form pyrite. This Mg 2+ was probably derived from connate water, circulating meteoric water, and from the alteration of Mg-bearing silicates (such as biotite) and carbonates. The sources of iron for formation of authigenic chlorites has been mostly detrital pigmentary hematite, biotite, magnetite, and Fe-Ti oxides.

Pyrite associated with sand-sized biotite and chlorite

Microprobe analyses of grains containing abundant pyrite suggests that iron was released during the alteration of biotite into illite (and /or kaolinite) and chlorite. In the case of "pyritized" chlorite grains, it is suggested that substitution of Fe 2÷ by Mg 2+ has occurred during chloritization of the parent biotite grains.

In sandstones containing abundant "pyritized" biotite grains (up to about 8% of the rock), pyrite occurs within interstitial chlorite as few scattered crystals, but never as extensive cement. This suggests that "pyritization" of biotite probably predates "pyritization" of interstitial chlorite and that iron released during illitization and chloritization of biotite was the earliest iron available for reaction with, and possibly complete utilization of, sulfide ions in these sandstones (cf. Berner, 1981). Illiti- zation of biotite has presumably occurred initially while the sediment was still subjected to oxic conditions, during a very early stage of burial. This also means that the iron released during this tri- to dioctahedral mica transformation has originally crystallized as fine pigmental iron oxides and /o r hydroxides. Subse- quently pyrite has probably been formed by the reaction of sulfide ions with these iron oxides and hydroxides (cf. Rickard, 1974), upon the burial of the sediments in the anoxic-sulfidic environment (Fig. 6). Part of the iron released from the earlier illitization of biotite (which initially crystallized as Fe-oxides and /o r hydroxides) presumably accounts for some of the interstitial pyrite. The iron oxides (and /or

190

E n v i r o n m e n t s B IOTITE B I O T I T E BIOTITE

Oxic ( b i o t i t e + i l L i t e + F e 2 0 3 ) ~ \ / i

A n o x i c - nonsulfid]c ( p o s t - o x i c )

Anoxic - sutf ldic

( b i o t i t e + i t t i t e + Fe 2+ ) ( b i o t i t e +chLor i te + K++T i ¢+) ( b i o t i t e + chLor i te + K++Ti z'÷)

./ ÷

(3

F i n a l m i n e r a l

a s s e m b t o g e

I tL i te + r e l i c biotite + p y r i t e + q u a r t s + m ino r a m o u n t s o~ b e m o t i t e

Fe- r lch ch to r l t e + t i t o n l u m oxides + q u a r t z + r e l i c biotite + s m a L L a m o u n t s o f py r i t e 0nd i t t i t e

F e M g - c h L o r i t e py r i t e + t i t a n i u m ox ides + q u a r t z + re l i c b io t l t e + smaLL a m o u n t s o f i LL i te

Fig. 6. A model for the alteration and "pyritization" of detrital biotite. For further explanation, see text

and Fig. 5.

hydroxide) might dissolve through direct or bacterial oxidation of organic matter (Irwin, 1980):

4FeOOH + CH 2 + H20 ~ 4Fe 2÷ + H C O ; + 7OH-

The above reaction supplies Fe 2+ (which could then react with sulfide ions to form pyrite) and HCO 3 necessary for the precipitation of calcite as the late authigenic phase in the sandstones studied. Chloritization rather than illitization of biotite has possibly occurred by rapid burial of sediments to depths below the oxic environment. Therefore, subsequent rapid decrease in redox potential prevailed and the iron released from alteration of biotite (and from dissolution of iron oxides) became increasingly soluble. The reaction of biotite into illite (analyses 1-3, Table 3) probably ceased due to an increase of ave2+ + aMg2+/aH+ ratio relative to aK+/aH+ ratio which instead of illitization caused chloritization of the biotite. When the chloritization of biotite occurred under conditions of high Mg 2+ activity (Mg substituted for octahedral FEZ+), pyrite was formed by the reaction of sulfide ions with excess Fe released from the biotite. This simplified model for alteration and pyritization of biotite is summarized in Fig. 6. There is no doubt that similar processes occurring in nature are more complicated.

Finally, the formation of discrete euhedral crystals, rather than framboidal pyrite, in the sandstones studied might reflect their relatively late-diagenetic origin (cf. Kalliokoski, 1966; Love and Amstutz, 1966; Ostwald and England, 1979; Raiswell and Plant, 1980). Discrete euhedral pyrite apparently forms under conditions of relatively lower rates of Fe z+ supply and /o r sulfate reduction (cf. Raiswell, 1982), than framboidal pyrite forms.

191

ACKNOWLEDGEMENTS

I a m gra te fu l to the s ta f f of the S E M L a b o r a t o r y , the A g r i c u l t u r a l U n i v e r s i t y of

N o r w a y , fo r the i r gene rous c o o p e r a t i o n . T h e m i c r o p r o b e ana lyses were c o n d u c t e d

at the G e o l o g i c a l M u s e u m , Oslo . C o n s t r u c t i v e c o m m e n t s a n d sugges t ions p r o v i d e d

by A .A. A 1 D a h a n , W.J . H a r r i s o n , P. Jo rgensen , S. N y b a k k e n , J.P. N y s t u e n , R .F .

Ra iswel l , K. Stanley, and two a n o n y m o u s rev iewers are g rea t ly app rec i a t ed . K.

G l o e r s e n k i n d l y t yped the m a n u s c r i p t and A. K a l j u s a a r d r a f t ed the f igures. F i n a n -

cial suppo r t was p r o v i d e d by the Swed i sh N a t u r a l Sc ience R e s e a r c h C o u n c i l ( N F R ) .

REFERENCES

AIDahan, A.A. and Morad, S., 1986a. Mineralogy and chemistry of diagenetic clay minerals in Proterozoic sandstones from Sweden. Am. J. Sci., 286: 51-80.

AIDahan, A.A. and Morad, S., 1986b. Chemistry of biotite and phyllosilicate intergrowths in sandstones. Clay Clay Miner., in press.

Barshad, I. and Kishk, J.M., 1968. Oxidation of ferrous iron in vermiculite and biotite alters fixation and replacability of potassium. Science, 162: 1401-1402.

Berner, R.A., 1970. Sedimentary pyrite formation. Am. J. Sci., 268: 1-23. Berner, R.A., 1981. A new geochemical classification of sedimentary environments. J. Sediment. Petrol.,

51 : 359-365. Bjorlykke, K., 1984. Formation of secondary porosity: How important is it? In: A. McDonald and R.C.

Surdam (Editors), Clastic Diagenesis. Mem. Am. Assoc. Pet. Geol., 37: 277-286. Blackbourn, G.A., 1984. Diagenetic history and reservoir quality of a Brent sand sequence. Clay Miner.,

19: 377-389. Drever, J.E., 1971. Magnesium-iron replacement in clay minerals in anoxic marine sediments. Science,

172: 1334-1336. Farmer, V.C., Russel, J.D., McHardy, W.T., Newman, A.D.C., Ahlrichs, J.L. and Rimsaite, J.H.Y., 1971.

Evidence for loss of protons and oxidized iron from oxidized biotites and vermiculites. Mineral. Mag., 38: 120-137.

Froelich, P.N., Klinkhammer, G.P., Bender, M.L., Luedtke, N.A., Heath, G.R., Cullen, D., Dauphin, P., Hammond, D., Hartmen, B. and Maynard, V., 1979. Early oxidation of organic matter in pelagic sediments of the eastern equatorial Atlantic: suboxic diagenesis. Geochim. Cosmochim. Acta, 43: 1075-1090.

Gilkes, P.J. and Suddhiprakarn, A., 1979. Biotite alteration in deeply weathered granite. I. Morpho- logical, mineralogical and chemical properties. Clays Clay Miner., 27: 349-360.

Gilkes, P.J., Young, R.C. and Quirk, J.P., 1972. The oxidation of octahedral iron in biotite. Clays Clay Miner., 20: 303-315.

Goodman, B.A. and Wilson, M.J., 1973. A study of the weathering of biotite using MSssbauer effect. Mineral. Mag., 39: 448-454.

Grossman, R.H., Leibling, R.S. and Scherp, H.S., 1979. Chlorite and its relationship to pyritization in anoxic marine environments. J. Sediment. Petrol., 49: 611-614.

Helgeson, H.C., Delany, J.M., Nesbitt, H.W. and Bird, D.K., 1978. Summary and critique of the thermodynamic properties of rock-forming minerals. Am. J. Sci., 278A: 1-229.

Howarth, R.W., 1979. Pyrite: its rapid formation in a salt marsh and its importance in ecosystem metabolism. Science, 203: 49-51.

Huggett, J.M., 1984. Controls on mineral authigenesis in coal measures sandstones of the east Midlands, U.K. Clays Clay Miner., 19: 343-357.

192

Irwin, H., 1980. Early diagenetic carbonate precipitation and pore-fluid migration in the Kimmeridge Clay of Dorset, England. Sedimentology, 27: 577-591.

Kalliokoski, J., 1966. Diagenetic pyritization in three sedimentary rocks. Econ. Geol., 61: 372-385. Li, Y.H., Bischoff, J. and Mathiue, G., 1969. The migration of manganese in the Arctic Basin sediments.

Earth Planet. Sci. Lett., 7: 265-270. Love, L.G. and Amstutz, G.C., 1966. Review of microscopic pyrite. Fortschr. Mineral., 43: 273-309. Lynn, D.C. and Bonatti, E., 1965. Mobility of manganese in diagenesis of deep-sea sediments. Mar.

Geol., 3: 457-474. Morad, S., 1983. Diagenesis and geochemistry of the Visings6 Group (Upper Proterozoic), southern

Sweden: a clue to the origin of color differentiation. J. Sediment. Petrol., 53: 51-65. Morad, S., 1984. Diagenetic matrix in Proterozoic graywackes from Sweden. J. Sediment. Petrol., 54:

1157-1168. Morad, S., 1986a. Albitization of K-feldspar in Proterozoic arkoses and graywackes from southern

Sweden. N. Jahrb. Mineral., Monatsch., 4: 145-156. Morad, S., 1986b. Mica-chlorite intergrowths in very low-grade metamorphosed sedimentary rocks from

Norway. N. Jahrb. Mineral., in press. Morad, S. and A1Dahan, A.A., 1986a. Discussion and comments on the paper: Electron-optical studies of

phyllosilicate intergrowths in sedimentary and metamorphic rocks. Mineral. Mag., 50: 340-343. Morad, S. and A1Dahan, A.A., 198619. Alteration of Fe-Ti oxides in sedimentary rocks. Geol. Soc. Am.

Bull., 97: 567-578. Ostwald, J. and England, B.M., 1979. The relationship between euhedral and framboidal pyrite in

base-metal sulphide ores. Mineral. Mag., 43: 297-300. Parry, W.T. and Downey, L.M., 1982. Geochemistry of hydrot19ermal chlorite replacing igneous biotite.

Clays Clay Miner., 30: 81-90. Peveraro, R.C.A. and Russell, K.J., 1984. Interpretation of wireline log and core data from a Mid-Jurassic

sand/shale sequence. Clay Miner., 19: 483-505. Raiswell, R., 1982. Pyrite texture, isotopic composition and the availability of iron. Am. J. Sci., 282:

1244-1263. Raiswell, R. and Plant, J., 1980. The incorporation of trace elements into pyrite during diagenesis of

black shales. Econ. Geol., 75: 684-699. Richard, D.T., 1974. Kinetics and mechanisms of the sulfidation of goethite. Am. J. Sci., 275: 636-652. Siever, R. and Kastner, M., 1972. Shale petrology by electron microprobe: Pyrite-chlorite relations. J.

Sediment Petrol., 42: 350-355. Veblen, D.R. and Ferry, J.M., 1983. A TEM study of the biotite-chlorite reaction and comparison with

petrologic observations. Am. Mineral., 68: 1160-1168. White, S.H., Huggett, J.M. and Shaw, H.F., 1985. Electron-optical studies of phyllosilicate intergrowths

in sedimentary and metamorphic rocks. Mineral. Mag., 49: 413-423. Wilson, M.D. and Pittman, E.D., 1977. Authigenic clays in sandstones: recognition and influence on

reservoir properties and paleoenvironmental analysis. J. Sediment. Petrol., 47: 3-31. Vidal, G., 1976. Late Precambrian microfossils from the Visingst~ beds in south Sweden. Fossils Strata, 9:

57 pp.

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Pyrite-chlorite and pyrite-biotite relations in sandstones