Carbonate Cementation in Sandstones: Distribution Patterns and Geochemical Evolution (IAS Special Publication 26)

CARBONATE CEMENTATION IN SANDSTONES

Carbonate Cementation in Sandstones Edited by Sadoon Morad © 1998 The International Association of Sedimentologists ISBN: 978-0-632-04777-2

SPECIAL PUBLICATION NUMBER 26 OF THE INTERNATIONAL ASSOCIATION OF SEDIMENTOLOGISTS

Carbonate Cementation in Sandstones DISTRIBUTION PATTERNS AND GEOCHEMICAL EVOLUTION

EDITED BY SADOON MORAD

b

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Carbonate cementation in sandstones/ edited by Sadoon Morad. p.

·em. - (Special publication

number 26 of the International Association of Sedimentologists) Includes bibliographical references and index. ISBN 0-632-0497 5-8 I. Sandstone.

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3. Rocks, Carbonate. I. Morad. Sadoon. II. Series: Special publication .. . of the International Association of Sedimentologists: no. 26. 1998

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Contents

1x

Preface Carbonate cementation in sandstones: distribution patterns and geochemical evolution

S. Morad

27

Origin and spatial distribution of early vadose and phreatic calcite cements in the Zia Formation, Albuquerque Basin, New Mexico, USA

J.R. Beckner and P.S. Mozley

53

Carbonate diagenesis and porosity evolution in sheet-flood sandstones: evidence from the Middle and Lower Lunde Members (Triassic) in the Snorre Field, Norwegian North Sea

S. Morad, L.F. de Ros, J.P. Nystuen and M. Bergan

87

Carbonate diagenesis in non-marine foreland sandstones at the western edge of the Alleghanian overthrust belt, southern Appalachians

KL. Milliken

107

Palaeogeographical, palaeoclimatic and burial history controls on the diagenetic evolution of reservoir sandstones: evidence from the Lower Cretaceous Serraria sandstones in the Sergipe-Alagoas Basin, NE Brazil

A.J. V Garcia, S. Morad, L.F. de Ros and I.S. Al-Aasm

141

Carbonate cements in the Tertiary sandstones of the Swiss Molasse basin: relevance to palaeohydrodynamic reconstruction

J. Matyas

163

Carbonate cement in the Triassic Chaunoy Formation of the Paris Basin: distribution and effect on flow properties

R.H. Worden and J.M Matray

179

Calcite cement in shallow marine sandstones: growth mechanisms and geometry

0. Walderhaug and P.A. BjtJrkum

v

Contents

vi

193

Origin of low-permeability calcite-cemented lenses in shallow marine sandstones and CaC03 cementation mechanisms: an example from the Lower Jurassic Luxemburg Sandstone, Luxemburg

N Molenaar 213

Geochemical history of calcite precipitation in Tertiary sandstones, northern Apennines, Italy

K L. Milliken, E.F. McBride, W Cavazza, U. Cibin, D. Fontana, MD. Picard and G. G. Zuffa

241

Diagenetic evolution of synorogenic hybrid and lithic arenites (Miocene), northern Apennines, Italy

E. Spadafora, L.F. de Ros, G. G. Zuffa, S. Morad and I.S. Al-Aasm

261

Carbonate cementation in Tertiary sandstones, San Joaquin basin, California

J.R. Boles

285

Carbonate cementation i n the Middle Jurassic Oseberg reservoir sandstone, Oseberg field, Norway: a case of deep burial-high temperature poikilotopic calcite

J.-P. Girard

309

Origin and timing of carbonate cementation of the Namorado Sandstone (Cretaceous), Albacora Field, Brazil: implications for oil recovery

R.S. de Souza and C.M. de Assis Silva 327

Structural controls on seismic-scale carbonate cementation in hydrocarbon-bearing Jurassic fluvial and marine sandstones from Australia: a comparison

J. Schulz-Rojahn, S. Ryan-Grigor and A. Anderson

363

Carbonate cementation-the key to reservoir properties of four sandstone levels (Cretaceous) in the Hibernia Oilfield, Jeanne d'Arc Basin, Newfoundland, Canada

R. Hesse and I.A. Abid

39 5

The significance of 813 C of carbonate cements in reservoir sandstones: a regional perspective from the Jurassic of the northern North Sea

C.!. Macaulay, A.E. Fallick, OM McLaughlin, R.S. Haszeldine and MJ. Pearson

409

Origin and significance of fracture-related dolomite in porous sandstones: an example from the Carboniferous of County Antrim, Northern Ireland

R. Evans, J.P. Hendry, J. Parnell and R.M Kalin

Contents

437

VII

Saddle (baroque) dolomite in carbonates and sandstones: a reappraisal of a burial-diagenetic concept

C. Spot! and J.K Pitman

461

Application of quantitative back-scattered electron image analysis in isotope interpretation of siderite cement: Tirrawarra Sandstone, Cooper basin, Australia

MR. Rezaee and J.P. Schulz-Rojahn

483

Carbonate cement dissolution during a cyclic C02 enhanced oil recovery treatment

L.K Smith

501

Index

Preface

Most special publications are proceedings of meet­

cementation and diagenetic evolution in oil-field

ings, and none covers specific topics of siliciclastic

sandstones from USA, North Sea, Brazil, Australia

diagenesis. It was,

and Canada. Chapter 17 evaluates the large-scale

therefore,

decided to invite

recognized experts from academia and industry to

carbon isotopic signatures in Jurassic sandstones

contribute to this lAS special publication. Each

from 13 North Sea oil fields. Chapter 18 discusses

manuscript was examined by two independent

fracture-related

referees. This has resulted in volume that contains

whereas Chapter 19 presents a reappraisal of the

dolomite in porous sandstones,

papers covering fairly broad aspects of carbonate

significance of saddle dolomite as an indicator of

cementation in sandstones in terms of the deposi­

burial diagenetic conditions in sandstones and car­

tional, tectonic and diagenetic settings of the basins

bonate rocks. Chapter 20 demonstrates the use of

studied. After my own opening review (Chapter 1),

quantitative back-scattered electron image analysis

contributions are arranged in the following order.

in the interpretation of the isotopic signatures of

Chapters 2-7, which deal with carbonate cementa­

carbonate cements in sandstones. The closing chap­

tion in continental sandstones, are followed by

ter discusses the dissolution of carbonate cement by

others (Chapters 8-11) dealing with cementation in

cyclic C02 enhanced oil recovery. S. Morad

marine sediments. Chapters 12-16 cover carbonate

IX

Spec. Pubis int. Ass. Sediment. ( 1 998) 26, 1 -26

Carbonate cementation in sandstones: distribution patterns and geochemical evolution S. M O R A D Sedimentary Geology Research Group, Institute o fEarth Sciences, Uppsala University, S-752 36 Uppsala, Sweden, e-mail sadoon. morad@geo. uu.se

ABSTRACT

Carbonate cements in sandstones are dominated by calcite, dolomite, ankerite and siderite, whereas magnesite and rhodochrosite are rare. The distribution patterns, mineralogy and elemental/isotopic compositions of carbonate cements vary widely, both temporally and spatially. The most important factors controlling these parameters during near-surface eodiagenesis include the depositional setting (e.g. rate of deposition, pore water composition, hydrogeology, climate, latitude and sea-level fluctuation), the organic matter content and the texture and detrital composition of the host sediments. During burial (mesodiagenesis) the important controlling factors include the temperature, residence time, chemistry and flow rates/pattern of subsurface waters, and the distribution patterns of eogenetic carbonate cements. As a result of mass balance constraints, burial carbonates are thought to be formed by the dissolution-reprecipitation (i.e. redistribution) of eogenetic carbonate cements and detrital carbonates. However, cements may also be derived internally from the dissolution of carbonate bioclasts, volcaniclastic material and calcium plagioclase, or externally from associated carbonate rocks, evaporites and mudstones. During uplift and erosion, carbonate cements are subjected to telogenetic alteration and dissolution. The imprints of eogenetic, mesogenetic and telogenetic conditions might be unequivocally reflected in the mineralogy and geochemistry of carbonate cements. However, eogenetic carbonates, particularly calcite and dolomite, may be subjected to recrystallization and resetting of isotopic signatures, fluid inclusion thermometries and elemental compositions.

INTRODUCTION

Carbonates are among the predominant cements in sandstones and thus an understanding of their distribution patterns and geochemical evolution is relevant to reservoir evaluation. Thorough studies of the composition and origin of carbonate cements in sandstones using modern analytical techniques have attracted sedimentary petrologists only in the past two decades. A proper study of carbonate cementation should be carried out within the dia­ genetic context of the host sandstones and should be based on as many analytical methods and as many background data about the sedimentary basin as possible. For instance, the timing and tempera­ ture of carbonate precipitation should not be de­ rived exclusively from thermometric measurements of fluid inclusions because inclusions may reCarbonate Cementation in Sandstones Edited by Sadoon Morad © 1998 The International Association of Sedimentologists ISBN: 978-0-632-04777-2

equilibrate subsequent to entrapment and give anomalously high temperatures. Thus the measured temperatures should be critically examined and cross-checked against petrographic observations, geochemical data on the carbonate and related cements, and the thermal history of the basins. Factors that control the geochemistry, abundance and distribution of carbonate cements are of prime importance in the understanding and prediction of porosity-permeability variations and in tracing the geochemical evolution of pore waters during the burial of sandstones and associated sediments. Moreover, the stable isotopic composition of near­ surface, eogenetic carbonates (e.g. in soil profiles) provides important clues to the palaeoclimatic con­ ditions (e.g. Ceding, 1984).

2

S. Morad

Water composition and flow pattern are of prime importance in determining the distribution and geochemical evolution of carbonate cements. These water properties vary considerably between near­ surface to shallow eodiagenesis and deep mesodia­ genesis. During eodiagenesis, the pore water chem­ istry is strongly controlled by the composition of the depositional waters, climate, detrital mineral com­ position and hydrology of the basin. Compared with eodiagenesis, water migration in the deep basinal regimes is limited by the decrease in poros­ ity and permeability of sandstones and associated rocks. The amounts and distribution patterns of mesogenetic carbonates, and hence the porosity­ permeability of the host sediments, are strongly constrained by the chemistry as well as timing, rate and extent of cross-formational water flow. Carbonate cements either indirectly enhance or deteriorate the reservoir properties of sandstones. Enhancement of reservoir properties occurs when (i) appreciable volumes of carbonate cements are dissolved, causing the formation of secondary po­ rosity and (ii) small amounts of carbonate cement are evenly distributed in the sandstones to support the overburden weight and prevent the collapse of framework grains and consequent elimination of primary porosity. Souza et al. ( 1995) demonstrated that a few per cent of dolomite cement is sufficient to prevent the collapse of Aptian reservoir sand­ stones from Brazil despite the high content of ductile lithic fragments. The deterioration of reservoir properties occurs when sandstones are massively cemented by car­ bonates. Although carbonate-cemented horizons are thin (,;;;; c. 2 m) and form only a minor portion of sandstone sequences, they may compartmentalize reservoirs by acting as barriers to water (and hydro­ carbon) flow both during migration from the source rocks to the reservoirs and during production (Kan­ torowicz et al., 1987; Carvalho et al., 1995). Com­ paction of sandstone sequences containing zones of laterally continuous carbonate-cemented horizons may lead to the development of overpressure in underlying, weakly cemented zones. Laterally ex­ tensive carbonate-cemented sandstones occur both in marine (Kantorowicz et al., 1987) and continen­ tal sequences (Arakel & McConchie, 1982). The chemical composition and distribution pattern of carbonate cements also has important implications for secondary oil recovery. For instance, ferroan carbonate reacts with injected acids to precipitate iron oxides/oxyhydroxides along the pore throats of

sandstones, causing a deterioration in permeability and oil recovery. The aim of this paper is to discuss the following topics: (i) the geochemical conditions of carbonate cementation in terms of organic-inorganic interac­ tions; (ii) the petrological and geochemical charac­ teristics of facies-related carbonate cements; (iii) the dissolution, recrystallization and replacement of carbonate cements during progressive sediment burial; and (iv) water-carbonate equilibrium states in some reservoir sandstones and deep-sea sedi­ ments on which pore water analyses and mineralog­ ical data are available.

GEOCHEMICAL ZONES OF CARBONATE CEMENTATION

Pore waters below the depositional surface undergo systematic changes in chemical and isotopic compo­ sitions. These changes occur within zones which are related to the availability of metabolizable organic matter, Fe- and Mn-oxides/oxyhydroxides, alkalin­ 2 ity and the concentration of dissolved 02 and so4 (Curtis, 1967, 1987; Claypool & Kaplan, 1974; Froelich et al., 1979; Berner, 198 1; Coleman & Raiswell, 1993). These geochemical changes (Fig. 1) are likely to be imprinted in diagenetic carbonates to an extent that recognition of the particular zone within which they precipitated is possible. As sand­ stones are relatively poor in organic matter, it is likely that the cementation related to the reactions discussed in the following section occurs partly in associated organic-rich mud. Oxic carbonates

Pore waters in oxic zones are characterized by a dissolved oxygen content greater than � 0.5 mill. Oxic carbonates prevail in: (i) subaerial environ­ ments, such as the vadose zone where the pores are periodically filled with gas, air and/or water; (ii) immediately below the sediment-water interface in aquatic environments; and (iii) in the phreatic zone below the water table where all the pores are regularly filled with water. The thickness of the oxic zone depends on the penetration, by diffusion or advection, of oxygen below the sediment surface. Oxygen diffusion into pore waters is largely con­ trolled by the organic content and the rate of deposition. In marine and lacustrine sediments the

CH20 + HN03 --'»- C02 + N2+ H20 [02] s::: 0.5 mill

Mn-Fe rich calcite and dolomite

CH20 + Mn4+ --'»- Mn 2+aq + C02 rhodochrosite (613Cmarin• � -6 %.) · CH20 +�pi·� HS"+ co; � F.e3+ � Fe2+aq . ,, ·

F�pobr calclte:andldolomlte '(1113c·:5>.- .2olto -10%o) ;.

�

g.

� §' -.. '1>



�

i;; � �

1:;'

Fig. 1. The geochemical zones of organic-inorganic interactions encountered during progressive burial of marine and continental siliciclastic sediments in

. various depositional settings. The reactions are not balanced and aim to show the main reactants and products. These zones include: (i) oxic (OX); (ii) suboxic which is composed of nitrate reduction (NR), manganese reduction (MnR) and iron reduction (FeR) subzones; (iii) bacterial sulphate reduction (SR); (iv) microbial methanogenesis (Me); and (v) thermal decarboxylation of organic matter (D). The authigenic carbonates characteristic for each zone and their o13Cp06 values are provided. Mg-siderite and Fe-magnesite are the more typical ferroan carbonates for burial diagenesis at elevated temperatures. Factors controlling anoxity of the bottom waters, and hence the sediments below the water-sediment interface in semi-closed and open marine (left) and in lacustrine (right) basins are illustrated too. Upwelling ofnutrient-rich waters (lower left) causes an increase in primary productivity, and hence higher organic matter content in bottom sediments (black). However, some of the organic matter may be derived terrestrially. High organic matter content in such open-marine sediments may lead to suboxic pore water compositions below the sediment-water interface. Anoxic non-sulphidic conditions in pore waters immediately below the sediment-water interface in lacustrine environment can be enhanced by rapid rate of organic matter accumulation (lower right). See text for further explanation.

w

4

S. Morad

concentration of dissolved oxygen in pore waters, and thus the thickness of the oxic zone, also de­ pends on the concentration of dissolved oxygen in bottom waters and the extent of bioturbation. Under oxic conditions, Mn- and Fe-oxyhydr­ oxides/oxides are stable and occur as discrete phases or are adsorbed onto the surfaces of other minerals such as clays. Therefore oxic carbonate cements have low Mn and Fe contents and are typical of near-surface, continental sediments with a very low organic matter content. In these sedi­ ments dissolved carbon is derived from the decay of plant remains in soil horizons and from atmo­ spheric C02 (Cerling, 1984). The 813C values of authigenic carbonates forming in vadose and shal­ low phreatic zones mostly vary between -I Oo/oo and -3o/oo, reflecting mixed sources of dissolved carbon derivation from the decay of c and c4 plants and 3 from atmospheric C02. In continental settings the 8180 composition of meteoric waters, and hence of carbonate cements, is strongly controlled by lati­ tude and climatic conditions (Suchecki et al., 1988; Morad et al., 1995). Marine oxic carbonates precip­ itate in open diagenetic systems and thus have 813C and 8180 compositions similar to those of unmod­ ified sea water. However, considerable variations in oxygen isotopic values occur due to variations in bottom temperature. Suboxic carbonates

When pore waters in both marine and continental sediments become significantly depleted in dis­ solved oxygen ( < 0.5 mill), three geochemical sub­ zones successively prevail (Fig. I ): (i) nitrate reduction into nitrogen (NR); (ii) manganese reduc­ 2 tion to Mn +aw (MnR); and, subsequently, (iii) 2 iron reduction to Fe + aq· (FeR). The type and elemental composition of carbonate cement formed are hence strongly controlled by the amount of Fe­ and Mn-oxides/oxyhydroxides. An increase in carbonate alkalinity in the NR subzone enhances the precipitation of carbonate cements with 8180 compositions similar to oxic carbonates, but with a slight enrichment in Mn and Fe and depletion in 13C. Rhodochrosite and siderite precipitate in the MnR and FeR subzones of sedi­ ments containing large amounts of Mn- and Fe­ oxides, respectively. Because the three subzones overlap, it is common to observe, such as in deep-sea sediments, that suboxic siderites and rhodochrosite are enriched in Mn and Fe, respec-

tively (Chow et al., 1996). Separation of the sub­ zones occurs, however, in some settings of the deep sea with very low sedimentation rates and a rela­ tively low organic content (Froelich et al., 1979). As in the oxic zone, the 813C values of suboxic carbonates in continental environments are con­ trolled by the 813C of atmospheric carbon and by the oxidation of terrestrial organic matter in the soil profile, whereas the 8180 values are mainly con­ trolled by latitude and climatic conditions. The 813C values of suboxic marine carbonates are influenced by carbon derived from sea water and from the oxidation of organic matter. The extent of 2 1 C incorporation into the carbonates depends on the amount and reactivity of the organic matter, the depth of the suboxic zone below the seafloor and the degree of bioturbation. The resultant 813C of dis­ solved carbon in the suboxic zone is �-6o/oo (McArthur et al., 1986). Carbonates from bacterial sulphate reduction

This process is most important in marine sediments where the pore waters contain appreciable amounts of dissolved sulphate. Bacterial sulphate reduction (BSR) operates when the pore waters are devoid of dissolved oxygen (i.e. anoxic). In euxinic basins the sediment experiences BSR diagenesis directly at the sediment-water interface (Fig. I); in other words, no oxic and suboxic phases are encountered (Curtis, 1987). Sulphate reduction is aided by anaerobic bacteria, as follows:

(I) It is uncertain whether this reaction enhances car­ bonate cementation. Conversely, in the presence of reactive iron, the precipitation of Fe-sulphide and a considerable increase in alkalinity occur as follows: 2 4FeOOH + 4S04 - + 9CH20 goethite =

4FeS + 9HC0 - +6H20 + H+ 3

mackinawite greigite

(2)

and 2 2Fe20 + 8S04 - + 15CH20 3 hematite =

4FeS2 + 15HC0 - +7H20 +OW 3 pyrite

(3)

The increase in alkalinity due to reactions (2) and (3) enhances carbonate precipitation in the BSR

Geochemical evolution of carbonate cements zone (Sholkovitz, 1973; Berner, 1984). Increased pore water alkalinity is recorded from organic-rich sediments which are influenced by BSR and pyrite formation (e.g. Berner et a/., 1970; Kastner et a/., 1990). Fez+ is incorporated into Fe-sulphides, thus cal­ cite and dolomite precipitating in the SR zone are largely Fe-poor. However, the amount of Fe that is incorporated into these carbonates depends on the amounts and reactivity of organic matter and detri­ tal Fe-minerals and the diffusion rate of sulphate from sea water. The latter is considerably influ­ enced by the degree of bioturbation, the sedimenta­ tion rate and the concentration of dissolved oxygen in bottom waters. Moreover, Coleman et a/. (1993) noted that some sulphate-reducing bacteria are capable of reducing Fe3+ to Fez+ using Hz and hence the availability of dissolved iron can be at least partly independent of the flux rates of sulphide ions. The decrease in concentration of sulphate due to reduction into sulphide is believed to en­ hance the precipitation of dolomite (Baker & Kast­ ner, 1981). Indeed, dolomite is common in organic­ rich sediments (Garrison et a/., 1984; Burns et a/., 1988; Slaughter & Hill, 1991; Baltzer et a/., 1994). In addition to ml!diating BSR, the oxidation of organic matter enhances dolomite formation by increasing the alkalinity and pH of the pore waters due to production of ammonia by the enzymatic degradation of protein (Slaughter & Hill, I 991). In marine sediments, the o13C signature of car­ bonate cements precipitated in the BSR zone is dominated by dissolved carbon derived from the oxidation of organic matter. However, mixing with carbon derived from the other sources such as marine pore waters and the dissolution of biogenic carbonates are also common. Generally, BSR is accomplished at shallow depths below the sedi­ ment-water interface to depths of a few hundred metres. Bacterial sulphate reduction diagenesis oc­ curs either homogeneously distributed in the sedi­ ments or locally in sediments undergoing overall oxic or suboxic diagenesis due to high local concen­ trations of organic matter, such as inside borings, burrows and bioclasts. Carbonates from microbial methanogenesis

This process prevails in anoxic marine and conti­ nental sediments and when sulphate is totally re­ duced in the BSR zone (Fig. I). Although the precise mechanism is poorly understood, methano-

5

genesis (Me) is believed to occur by the fermenta­ tion of simple organic compounds, e.g. acetate (4) or via Hz production and subsequent C02 reduc­ tion: (5) The overall reaction of microbial methanogenesis can be envisaged as follows: (6) Both reactions (4) and (5) probably occur in the Me zone. The o13C values of C02 derived from these reactions depend on the specific microbial process involved. Where reaction (4) dominates, such as in freshwater environments, C02 inherits the o13C of the acetate, typically 5-l Oo/oo heavier than bulk carbon in the precursor organic matter (o13C � -I Oo/oo to -25o/oo), whereas the methane inherits the o13C value (-55o/oo to -60o/oo) of the methyl groups (Galimov, 1985; Whiticar et a/., 1986; Clayton, 1994). Reaction (5), which dominates in marine sediments, involves a strong kinetic carbon isotopic 2 fractionation causing the enrichment of CH4 in 1 C (o13C � -75o/oo) and enrichment of C02 in 13C. Residual C02 due to progressive, but incomplete, reduction by H2 into methane attains o13C values up to about +21o/oo (Deuser, 1979). Therefore it appears that o13C values of C02 in the Me zone vary between about -25o/oo and +21o/oo (cf. Whiticar et a/., 1986). Regardless of the dominating Me pathway, the earliest formed methane is isotopically 2 more enriched in 1 C. High rates of C02 production by reaction (4), which cause no change in the pH of the pore waters, lead initially to the dissolution rather than precipitation of carbonates. Carbonates precipitated in this zone have 'intermediate' o13C values (mostly between -22 and +2o/oo). Conversely, carbonates that have very positive carbon isotopic values are relatively rare (cf. Clayton, 1994). As a result of the anoxic, low sulphate concentra­ tions in the Me zone, carbonates expected to form include siderite and ferroan dolomite/ankerite (Gautier & Claypool, 1984). The precipitation of these carbonates occurs in sediments rich in reac­ tive detrital iron (Coleman, 1985), as follows: (7) 2Fe20 + 7CH20 4FeC0 + 3CH4 + H20 3 3 The solubility of methane in pore waters is limited and depends on the pressure, temperature and salinity. Excess methane dissipates upwards and is =

S. Morad

6

oxidized anaerobically in the BSR zone and aerobi­ cally in the suboxic zones as follows: 2 CH4 + S04 -

(8) HS- + HC0 - + H20 3 (9) CH4 + 202 H20 + HC0 - + H+ 3 2 These two reactions contribute 1 C to the pore waters in the sulphate reduction and, particularly, the suboxic zones. Methane seepages on the seafloor are accompanied by the formation of authigenic 2 calcite and aragonite that are highly enriched in 1 C (Hovland et a!. , 1 987). Within the zone of methane oxidation, rates of sulphate reduction may be seasonally and spatially variable. Iron carbonates form in the BSR zone due 2 to reduction of Fe3+ to Fe + by sulphate-reducing bacteria (Coleman et a!., 1 993). Alternating zones of dolomite and siderite (Morad, unpublished data) occur due to fluctuations in the positions of the transition from FeR to BSR and from BSR to Me zones. Alternating bands of siderite (o13C � -6o/oo) and ankerite (o13C �- I Io/oo in Jurassic marine sandstones from the Barents Sea have probably been formed due to this FeR to BSR or BSR to Me fluctuation mechanism (Morad et a!. , 1 996). Fluc­ tuations in the geochemical zones are brought about due to the episodic oxygenation of anoxic basins or changes in the rate of sedimentation and flux of organic matter. In some cases, deep-sea carbonates have o180 values that cannot be explained even if the bottom water temperature is assumed to be o·c (Wada et al., 1 982). Such an anomalous 180-enrichment of carbonates (o180PoB up to +7.9o/oo) has been ar­ gued to be related to the destabilization of gas hydrates (Matsumoto, 1 989). The Me zone may extend from the surface to burial depths corresponding to a temperature in­ crease to about 7 5 ·c , where biological activity is decreased or largely inhibited. However, formation waters at temperatures> 8o·c with o13C values as high as + 5%o have been reported by Carothers & Kharaka (I 980), suggesting that methanogenesis may occur at higher temperatures. =

=

Carbonates from thermal decarboxylation of organic matter

As bacterial activity diminishes due to an increase in temperature, the diagenetic reactions in which organic matter plays an important part will be thermally controlled to temperatures perhaps as

high as � 2so·c (Carothers & Kharaka, 1 978, 1 980; Surdam e t al., 1 984; Giordano & Kharaka, 1 993). These workers have argued that there is sufficient evidence indicating that carboxylic acids, as well as C02 and H20, are produced in the early stages of the thermocatalytic degradation of ali­ phatic acids incorporated in kerogen before hydro­ carbon generation. At temperatures between 80 and 1 20"C relatively high concentrations (up to I 0 000 mg/ 1) of carboxylic acids, particularly ace­ tate, are detected in oil-field brines (Hanor & Workman, I 986; Kharaka et al. 1 986; MacGowan & Surdam, 1 990). Over this temperature range the pH of the carbonate system is externally buffered by carboxylic acid anions (Surdam et al. , 1 984). Hence the decarboxylation of organic matter and conse­ quent increase in Pco, would enhance the precipi­ tation rather than dissolution of carbonate cements. External pH buffering and enhanced carbonate precipitation may also occur due to silicate reac­ tions (e.g. the dissolution and albitization of detrital feldspar, chloritization of mica) in the diagenetic system (Smith & Ehrenberg, 1 989; Hutcheon & Abercrombie, 1 990). The o13C values of the carbon derived from organic matter is � - 1 5o/oo. The o 13C of carbonate cements in this zone is usually consid­ erably influenced by the redistribution of earlier formed carbonates, but is �- I Oo/oo. At temperatures greater than � 1 0o·c , thermal degradation of carboxylic acids produces methane and carbon dioxide (Surdam et al. , 1 9 84). As the carboxylic acid anions are consumed due to increas­ ing temperature, the carbonate system becomes internally buffered, and thus the pH may decrease due to increased Pco, in the system, leading to carbonate dissolution and the enhancement of sec­ ondary porosity (Surdam et al. , 1 984). Factors influencing the thermal destruction rate of organic acids include coupled sulphate reduction and hy­ drocarbon oxidation, and the mineralogy of host sediments (Bell, I 99 1 ); the presence of hematite causes rapid rates of acetic acid decomposition. Over the temperature interval 1 20- 1 60·c the carboxylic acid anions completely decarboxylate and the alkalinity is dominated by the carbonate system. Consequently, any increase in Pco,- will cause further dissolution. A variety of carbonate cements occurs in the de­ carboxylation zone depending on the mineralogy of the host sediments and earlier formed carbonates, as well as incursion by deep-seated thermobaric waters. Sediments containing abundant reactive, detrital Fe-minerals result in the formation of

Geochemical evolution of carbonate cements ferroan calcite and ankerite (e.g. Kantorowicz, 1 98 5 ).

FACIES-RELATED DISTRIBUTION OF CARBONATE CEMENTS

Like other diagenetic minerals in siliciclastic se­ quences, eogenetic carbonate cements may display a strong relationship with depositional facies in continental and marine settings. Continental calcite and dolomite

Calcretes and dolocretes are the dominant forms of carbonate cements in continental and nearshore sediments, which develop in warm to hot, arid to semi-arid regions, with low, seasonal rainfall and high evaporation (Goudie, 1 98 3). However, cal­ cretes composed of low-Mg calcite may develop in wet, cold areas and in dry Arctic soils by freezing (Swett, 1 9 74; Bunting & Christensen, 1 9 80; Drozdowski, 1 9 80). Strong et a!. ( 1 992) found that in cold, wet areas calcrete formation is enhanced by the presence of abundant carbonate clasts and a high degree of biological activity beneath forest covers. The stable isotopic composition of these carbonates is a powerful tool for inferring palaeo­ environmental and palaeo-ecological variables such as climate, vegetation type and atmospheric levels of C0 (Cerling, 1 984, 1 99 1 ; Cerling & Hay, 1 986; 2 Cerling et a!., 1 989; Mack et a!. , 1 99 1 ; Mora et a!., 1 99 1 ; Driese & Mora, 1 993). Carbonate precipitation in the vadose zone of hot arid to semi-arid regions is enhanced by a decrease in Pco, and PH,o due to increasing temperature and evaporation. Conversely, carbonate leaching is en­ hanced by a humid climate, which prevents the evaporative concentration of dissolved Ca2+ and Mg2+. Loss of water through uptake by plants was argued by Klappa ( 1 980) to be a likely mechanism for the precipitation of carbonates around roots. Carbonate precipitation around roots (rhizocre­ tions) may also be enhanced by microbial activities (Krumbein, 1 968) and an increase in alkalinity due to the decay of dead plants. Sources of Ca and Mg for calcrete and dolocrete are uncertain, but are often believed to be wind­ blown dust. Calcium and Mg may also be derived from pyroclastic material (Bestland & Retallack, 1 993) and oceanic aerosols (Quade et a!., 1 995). These sources are also relevant to phreatic carbon­ ates. In some cases the groundwater may bring ions

7

from carbonate rock terranes to siliciclastic se­ quences. Additional sources include Ca dissolved in rainwater ( � 6-7 ppm; Goudie, 1 973), Ca­ plagioclase, Ca in tissues of certain plants, and carbonate bioclasts (e.g. land snails). Dissolution of carbonate grains may occur as a consequence of: (i) a build up of Pco, in the vadose zone due to the extensive respiration of plants and micro-organisms; (ii) an increase in the concentra­ tion of organic (humic-fulvic) acids due to secre­ tion by, or decay of, plants; and (iii) mixing between waters with chemically different compositions, par­ ticularly in terms of Pco, (e.g. vadose and phreatic waters), which is referred to as mixing corrosion (Wigley & Plummer, 1 976). Calcretes and dolocretes occur as concretions and laterally extensive cements in floodplain and nearshore sediments (Figs 2 & 3). The carbonate­ cemented zones reach thicknesses exceeding I 0 m and dimensions of over I 0 km x I 00 km (Arakel & McConchie, 1 982; Arakel, 1 986). Calcretes and dolocretes may also develop in fluvial channel sandstones (Tandon & Narayan, 1 98 1 ; Arakel et a!., 1 990; Arakel, 1 99 1 ). These are dominantly phreatic carbonates formed by the dissolution and re­ precipitation of carbonate intraclasts derived from the erosion of floodplain pedogenic calcretes. The high permeability of channel sandstones also en­ hances the dissolution and kaolinization of detrital silicates, particularly in semi-arid regions with ac­ tive groundwater systems (Fig. 2). Dolocretes are common in fine-grained distal fluvial facies, whereas calcretes dominate in coarse­ grained, proximal facies (Fig. 3). The precipitation of dolomite is enhanced by an increase in the Mg/Ca ratio of flowing groundwaters due to the precipitation of calcite in proximal sediments and to evaporative ionic concentration. Dolomite pre­ cipitation in lacustrine environments is believed to occur from mixed groundwater and lake brines, which sink into the sediments during periods of intensive evaporation and density increase (Colson & Cojan, 1 996; Spot! & Wright, 1 992). Calcretes and dolocretes composed of alternating bands of calcite and ferroan to non-ferroan dolomite are believed to reflect precipitation from mixing be­ tween fresh phreatic waters and more saline, vadose waters (Watts, 1 980; Morad et a!., this volume; Saigal et a!., in preparation). Dolomite precipitates when the pore waters are enriched in Mg2+ due to its evaporative concentration, whereas calcite pre­ cipitates from fresh waters during rainy periods. Shallow marine sands and gravel rich in carbonate

8

S. Morad

Fig. 2 Distribution of eogenetic carbonates and clay minerals in a meandering fluvial system under semi-arid climatic conditions.

Fig. 3. Variations in relative importance of the different geochemical zones of diagenesis (see text) and the diagenetic

minerals formed in a profile covering proximal to distal continental arid to semi-arid environments, as well as subaquatic marine environments. Mn- and Fe-oxides should be encountered in the oxic zone of these different settings.

Geochemical evolution of carbonate cements bioclasts may also be incurred and cemented by meteoric waters. In some cases the cement is low-Mg calcite which occurs as concretions oriented parallel to the flow pathways of groundwater (Johnson, 1 989; McBride et a!., 1 994). The concretions may contain cracks that result from repeated wetting and drying events. These cracks are filled by clays and silt in areas which are episodically flooded, or filled by phreatic carbonate cement composed of coarsely crystalline calcite, dolomite or alternating bands of calcite/dolomite, or rarely fibrous radiaxial calcite (Saigal et a!., in preparation). Criteria for the identification of va­ dose cements include: (i) pendant or meniscus texture; (ii) carbonate precipitation in close relation to rootlets (rhizocretions); (iii) displacive and grain­ shattering carbonate cements (Braithwaite, 1 989; Saigal & Walton, 1 988); and (iv) patchy lumines­ cence due to episodic cementation related to tem­ poral filling of the pores with water. Calcretes and, particularly dolocretes in hot, arid climates are commonly associated with Mg-clays (sepiolite and palygorskite), silcrete and gypcrete (Watts, 1 9 80; El-Sayed et a!., 1 99 1 ; Spot! & Wright, 1 992; Colson & Cojan, 1 996). However, authigenic silica is preferentially associated with calcretes and dolocrete developed on chemically unstable volca­ nic bedrocks (Hay & Wiggins, 1 980). Conversely, carbonates formed under semi-arid conditions con­ tain both smectite and kaolinite (Fig. 2) (Morad et a!., this volume). Dolocretes are often closely asso­ ciated with ultramafic bedrocks, which result in an 2 2 increase in the Mg + /Ca + ratio of the groundwa­ ters (Watts, 1 9 80; Maizels, 1 9 87; Bums & Matter, 1 995). In some occurrences there is a close link of dolocrete formation with dolostone bedrock, such as in Miocene palaeosols from Spain (Alonso Zarza et a!., 1 992). Dolomite cement is also common in sandstones that are closely associated with evapor­ ite deposits in coastal and inland sabkha settings (Strong & Milowdowski, 1 987; Shew, 1 99 1 ; James, 1 992; Morad et al., 1 995). In these settings, dolo­ mite precipitation is enhanced by an increase in the Mg/Ca ratio of pore waters due to the evaporation of marine or mixed marine/meteoric waters (Patterson & Kinsman, I 982) and the precipitation of calcite and calcium sulphate cements (Kinsman, 1 969). Marine calcite and dolomite

Eogenetic calcite cement dominates in shallow marine siliciclastic sediments, and accompanies sulphate reduction and methane oxidation (Kantor-

9

owicz et a!., 1 987; Wilkinson, 1 99 1 ). Dolomite occurs in relatively small amounts, mainly in the sulphate reduction zone as overgrowths on detrital dolomite and by the diagenetic replacement of cal­ cite and aragonite precursors. The main sources of ions for carbonate cements are sea water, biogenic carbonates and carbonate intraclasts. Sea water Ca, Mg and HC0 - are introduced into the pore waters 3 by diffusion, or advection by storms and tidal cur­ rents. Chemical gradients are established due to the onset of carbonate precipitation as a consequence of the oxidation of local concentrations of organic matter, and hence an increase in alkalinity. Berner ( 1 968) demonstrated experimentally that the bacte­ rial decomposition of fish caused an increase in pH of the solution and consequently the precipitation 2 of Ca + as a mixture of calcium fatty acids salts or soaps. Berner ( 1 968) suggested that some ancient calcite concretions, especially those enclosing the skeletons of soft-bodied organisms, may have ini­ tially formed as calcium soaps which later con­ verted to CaC0 . No evidence of this process has 3 yet been provided for natural settings. The rapid (tens of years) carbonate cementation (high-Mg calcite and aragonite) of sand deposits which occurs in Recent tropical and subtropical marine coastal settings and results in the formation of tight beach rocks (Krumbein, 1 979; Amieux et a!., 1 989; Strasser et a!., 1 9 89; Guo & Friedman, 1 990) was probably also common in the geological past. Carbonate precipitation occurs in the marine vadose zone within intertidal and low supratidal sediments, most probably due to evaporation and C02 degassing (Hanor, 1 978) and photosynthesis by algae (Holail & Rashed, 1 992). There are no well-established criteria with which to recognize ancient beach rocks, as they are often subjected to recrystallization and dolomitization ( Ingvald, 1 995). However, dolomitized beach rocks usually preserve two characteristic features: (i) the presence of carbonate fringes around well rounded, unre­ placed framework grains, and (ii) the microcrystal­ line habit of the intergranular carbonate (see Fig. 4). Cement fabrics (see Fig. 4) typically comprise rims of numerous scalenohedral crystals or syntax­ ial overgrowths around carbonate bioclasts and intraclasts which grade into micritic or blocky crystals towards the pore centre (Spadafora et a/., this volume). The earliest formed rims and over­ growths are often non-luminescent due to a lack of Mn, indicating an oxic marine origin. Sands enriched in detrital carbonates and bioclasts are rapidly cemented by fringing calcite while on the

10

S. Morad

seafloor. This leads to stabilization of the arenite framework against porosity destruction by compac­ tion during subsequent burial. Early cementation is kinetically enhanced by nucleation on carbonate substrates. Upon burial, the progressive addition of coarse blocky or mosaic calcite over early calcite (Wilkinson, 1 99 1 ; Carvalho et a!., 1 995) and minor dolomite cements may lead to the formation of extensively cemented sandstones. These eogenetic cements are strata-bound, nodular or laterally con­ tinuous from hundreds of metres to several kilome­ tres (Kantorowicz et a!., 1 987; Prosser et a!., 1 993). Shallow marine sandstones often are enriched in biogenic carbonates which act as nuclei for calcite precipitation and as a cement source during burial (Bj0rkum & Walderhaug, 1 990). The deposition of such sandstones occurs in wave- and storm­ dominated, shallow marine environments, and to a smaller extent in muddy, fair weather sediments, tidal channels and tidal point bars. Shell-dominated layers also form by reworking into the slope apron (Hendry et a!., 1 996) and as a consequence of short­ term mortality due to catastrophic events such as an increase in water-column turbidity and a decrease in dissolved oxygen concentration. In warm, oxygen­ ated marine pore waters the bioclasts themselves usually do not dissolve because they are originally formed in equilibrium with sea water. However, the dissolution of metastable aragonite and high-Mg cal­ cite may begin in the suboxic and bacterial sulphate­ reduction zones (Morse & Mackenzie, 1 990). In nodular cemented sandstones, the areas left uncemented often reveal evidence of later burial diagenetic modifications, such as compaction and quartz cementation (Morad et a!., 1 995). Burial cements are believed to be sourced from meteoric or dissolution of detrital carbonates and bioclasts (cf. Wilkinson, 1 99 1 ). As the sandstone framework is expected to be stabilized due to early cementa­ tion, the burial dissolution of bioclasts may be recognized by oversized pores and mouldic pores filled with cement. Although abundant skeletal bioclastic fragments play an important part in the development of calcite-cemented sandstones, they should not be considered as the only source of such cements. Evidence for this is the common presence of calcite­ cemented sandstones in Precambrian sequences. Additional evidence is the absence of bioclastic carbonates in Jurassic sandstones with strata-bound calcite cements (Bj0rkum & Walderhaug, 1 993; Prosser et a!., 1 993). This suggests that other sources such as sea water and carbonate mud

intraclasts are at least as important as bioclasts. Highly reactive volcaniclastic sediments may also enhance carbonate cementation at shallow depths below the seafloor. Alteration of these sediments may cause the establishment of Ca2+, Mg2+ and HC03- diffusion gradients between pore waters and overlying seawater (Morad & De Ros, 1 994). The domination of calcite over other carbonates in volcaniclastic sediments (De Ros et a!., 1 996) is unclear, but may be related to the preferential incorporation of Fe2+ and Mg2+ in trioctahedral smectite, and to a diagenetically open system with respect to the overlying seawater. The mechanisms bringing marine pore waters into supersaturation with respect to calcite in volcaniclastic sediments are poorly understood. Another potential mechanism responsible for the formation of laterally continuous, strata-bound calcite-cemented sandstones is the episodic up­ welling of anoxic seawater (see Kempe, 1 990; Grot­ zinger & Knoll, 1 995). The upwelling of such high alkalinity waters to shelf and coastal areas may occur subsequent to periods of sea water stratifica­ tion accompanying sea-level rise. Subsurface, carbonate-cemented sandstone beds can be recognized from geophysical well logs and cores. Moreover, concretionary cemented sand­ stones are differentiated from continuously ce­ mented horizons based on these methods. The latter sandstones show as tight intervals on sonic, density and neutron logs. Scattered small concre­ tions give a less distinct response on density and neutron logs because of their limited lateral extent and show resistivity readings that vary around the borehole. Unlike eogenetic strata-bound cementa­ tion, continuous mesogenetic carbonate-cemented sandstone horizons are structurally controlled and cut across stratification when precipitation is re­ lated to water flow along faults. Calcite cement in ancient marine sediments is consistently a low-Mg variety (Magaritz et a!., 1 979; Spadafora et a!., this volume), which is either a primary precipitate or results from the stabilization of metastable high-Mg calcite and aragonite precur­ sors. There is evidence indicating that inorganic carbonates precipitated from sea water have varied between low-Mg calcite during periods of global warming (greenhouse mode due to an increase in atmospheric Pco ) and high-Mg calcite and arago­ nite during periods of global cooling (Sandberg, 1 983). Increased atmospheric Pco, has been related to periods of high plate tectonic activity, which leads to the release of more C0 derived from the 2

ll

Geochemical evolution of carbonate cements metamorphism of calcareous sediments at subduc­ tion zones (Wilkinson et al., 1985). Calcite stabili­ zation during these periods is further enhanced by lower Mg/Ca ratios in sea water due to the interac­ tion with ejected oceanic crust at mid-oceanic ridges. Thus low-Mg calcite fringes in some ancient marine sediments, such as the extensively studied Jurassic sandstones of the North Sea (Girard, this volume), are likely to be primary. This issue can be extended further to include a discussion on the variation in abundance of dolo­ mite in marine sandstones during geological times. This variation in the calcite/dolomite ratio, with the greater abundance of dolomite in old sedimentary rocks, is probably not only the result of burial diagenesis, but also due to changes in palaeo­ oceanographic conditions (Given & Wilkinson, 1987). Possible factors include the following(Purser et a!., 1994): (i) climate-tropical to subtropical climate favours the precipitation of dolomite (Tucker & Wright, 1990); and (ii) global sea-level change-sea-level rise leads to the incursion of nearshore areas by sea water, which, upon mixing with meteoric waters and evaporation, enhances the local precipitation of dolomite due to an increase in the Mg/Ca ratio of pore waters.

2 tremely 1 C-rich, isopachous Mg-rich calcite has also been reported from other modern non-tropical shallow marine terrigenous sediments, including the northeast USA shelf(Hathway & Degens, 1969), the Mississippi River delta (Roberts & Whelan, 1975) and the Kattegat Sea (J0rgensen, 1976, 1979). Simulations of marine-meteoric mixing (e.g. Plummer, 1975; Wigley & Plummer, 1976) pre­ dicted calcite oversaturation in waters with 20-70% sea water. However, the saturation degree of the mixed waters varies depending on the initial calcite saturation index, Pc02 and temperature. Neverthe­ less, predictive models constructed by Frank & Lohmann (1995) for low-Mg calcite precipitation in

Sea level1

Sea level2

Mixed marine-meteoric water carbonates

The degree of mixing between marine and meteoric waters, and hence the mineralogy, texture and pattern of carbonate cementation in coastal sand­ stones, are strongly influenced by sea-level fluctua­ tion (Fig. 4). The precipitation of eogenetic calcite and dolomite in nearshore sandstones occurs as alternating bands formed by precipitation from mixed marine-meteoric waters (Morad et a!., 1992). Evidence from present day settings suggests that the influence of marine mixing with fresh groundwater, and hence dolomite formation, may extend landward for distances of 25-30 km (Ma­ garitz et a!., 1981). Carbonate precipitation from mixed waters is enhanced by an increase in alkalinity due to the oxidation of organic matter and methane (Lunde­ gard, 1994). Gas pockets are common in Holocene sediments rich in organic matter (e.g. McMaster, 1984). Nelson & Lawrence (1984) and Simpson & Hutcheon (1995) reported the formation of Ho­ locene, high-Mg calcite nodules (1513C � -49o/oo to -7o/oo) in hybrid, bioclastic deposits of the modem Fraser River delta(� 49.N) due to methane oxida­ tion close to the seafloor. Early diagenetic, ex-

phreatic calcite vadose calclle

1.2 Meteoric water vadose and/or phreatic

meteoric caicite 3

4

Mixed to meteoric

5

h!IJh-Mg calcite/aragonite fnnges pores

Fig. 4 Influence of sea-level drop on the composition and texture of carbonate cements in sandstones situated in shallow marine, coastal and nearshore settings.

12

S. Morad

carbonate sediments from mixed waters suggest that the zone of oversaturation with respect to calcite can expand to encompass the full range of mixing. If high Mg/Ca ratios are maintained, arago­ nite rather than calcite may rarely precipitate from mixed waters (Kimbell & Humphrey, 1994). Siderite

Siderite precipitates from reducing, non-sulphidic pore waters that evolve in the suboxic and micro­ bial methanogenesis zones of all depositional envi­ ronments. These geochemical conditions occur in organic-rich sediments containing appreciable amounts of reactive iron minerals and in which the pore waters are so/--poor meteoric or brackish (Postma, 1982). Siderite is most common in continental and coastal sediments due to the much lower contents of dissolved sulphate in meteoric and brackish waters than in sea water. In these environments, small amounts of iron sulphide are formed, which allows an increase in Fe2+ concentration in pore waters, and hence promotes siderite formation. Siderite is abundant in fine-grained, organic-rich marsh and swamp sediments associated with deltaic and coastal sediments. Siderite slightly enriched in Ca and Mg formed in Holocene intertidal marsh and sandflat sediments from both marine and mixed marine-meteoric pore waters (Pye et al., 1990; Moore et al., 1992). In these sedimentary facies, siderite is closely associated with pyrite and Fe­ dolomite/ankerite. In fluvial sediments, siderite preferably forms in fine-grained floodplain and crevasse splay or in oxbow lake and pond sedi­ ments. The presence of plant remains in semi-arid to semi-humid regions enhances its formation (Fig. 3). Authigenic siderite spherules and thread­ like morphologies related to the replacement of detrital mica (Morad et al., this volume) are com­ mon in pedogenic profiles (Besly & Fielding, 1989; Kantorowicz, 1990; Browne & Kingston, 1993). According to Mozley ( 1989a), the elemental com­ position of siderite is controlled by the chemistry of depositional waters, with meteoric siderites being more enriched in Mn, but depleted in Ca and particularly Mg compared with siderite in marine sediments. However, Morad et al. (this volume) found that eogenetic siderites formed in a continen­ tal setting are highly enriched in Ca and Mg. Additionally, high-Mg siderites are typically formed at increased temperatures (Morad et al., 1994).

Thus, to apply the findings of Mozley ( 1989a), it is important to determine precisely the diagenetic regime of siderite formation. Unlike calcite and dolomite, siderite rarely forms as an extensive pore-filling cement, but rather as discrete fine crystals, spherules and nodules scat­ tered in the host sediments. Nevertheless, Baker et al. ( 1996) found that early diagenetic siderite con­ cretions (0.5-2 mm) form up to 30% of Triassic sandstones and mudstones from eastern Australia. Laterally continuous siderite-cemented offshore shelf sandstone sheets ( 15 em thick) occur in Upper Cretaceous sequences from Canada (McKay et al. , 1995). Rhodochrosite

Rhodochrosite occurs mainly in fine-grained ma­ rine and brackish water sedimentary basins, such as the Baltic Sea (Lynn & Bonatti, 1965; Suess, 1979; Pedersen & Price, 1982; Minoura, 199 1). In deep­ sea sediments, Ca- or Fe-rich rhodochrosite occurs as scattered crystals, microspherules and as nodules within host pelagic sediments (Coleman et al., 1982; Wada et al. , 1982; Matsumoto, 1992; Chow et al. , 1996). However, Bruhn ( 1993) observed fine-crystalline rhodochrosite nodules in fine­ grained sandstones and siltstones of Lower Tertiary, submarine turbidites from Brazil. Magnesite

Eogenetic magnesite cement in sandstones is rela­ tively rare because its formation requires pore waters to be enriched in Mg2+ and depleted in Ca2+, SO/- and Cl-. These conditions may occur in arid climates in which marine pore waters evap­ orate and become successively saturated with re­ spect to calcium carbonates, calcium sulphates and halite, such as in sabkha settings (Kinsman, 1969; Morad et al., 1995). Continental brines enriched in Mg2+ are also suitable for the formation of eoge­ netic magnesite due to the low sulphate and chlo­ ride ion concentrations. Most recent magnesite cements form in the fine-grained sediments of alkaline/saline lakes (Last, 1992; Warren, 1990) and, less commonly, in freshwater lacustrine sedi­ ments (Zachmann, 1989). Magnesite precipitates at depths of a few decime­ tres below the sediment-water interface, such as in the ephemeral salt pans of Recent playa lakes in north-east Spain, where precipitation is enhanced

Geochemical evolution of carbonate cements by increases in carbonate alkalinity due to bacterial activity (Pueyo Mur & Ingles Urpinell, 1 987). In the Permian Rotliegend reservoir sandstones from southern North Sea, intergranular, eogenetic mag­ nesite occurs in interdune sabkha as well as dune and fluvial facies (Purvis, 1 992). Eogenetic magne­ site occurs as nodules and layers in Permian playa lake mudstones and as intergranular cements in alluvial fan sandstones from Austria (Spot! & Bums, 1 994). The precipitation of this magnesite has been attributed to high-Mg brines derived from the weathering of Devonian dolostones and associ­ ated massive magnesite deposits in the catchment area (Spot! & Burns, 1 994). Marine eogenetic magnesite is also known to pre­ cipitate in deep-sea sediments. Matsumoto ( 1 992) described rhombic, microcrystalline (2- 1 5 J.Lm) Ca­ Mn-Fe rich magnesite and Fe-Mn rich lansfordite (hydrous Mg-carbonate) in Miocene to Pliocene mudstones from ODP Site 799 in the Japan Sea. He concluded that on progressive burial and increase in temperature (� 435 mbsf, T � 43 oq, the meta­ stable lansfordite is transformed into magnesite.

DISSOLUTION OF CARBONATE CEMENTS: MECHANISMS AND CONSEQUENCES

When carbonate cements are subjected to physico­ chemical conditions that vary considerably from those under which they formed, they may dissolve and re-precipitate at various scales. Carbonate dis­ solution and the creation of secondary porosity may occur during eodiagenesis or telodiagenesis or in response to progressive burial. Eogenetic secondary pores may survive subsequent burial and compac­ tion in sandstones that have been subjected to early overpressuring or hydrocarbon emplacement, or if dissolution is incomplete, and leave evenly distrib­ uted remnants of carbonate cement. The scales of carbonate redistribution, and thus reservoir quality enhancement, are difficult to con­ strain. Several workers have argued that the reser­ voir properties of sandstones are greatly enhanced due to large-scale carbonate dissolution (L0n0y et a!., 1 986; Schmidt & McDonald, 1 979). As the un­ dersaturated waters have to circulate through large volumes of permeable sediment to cause economi­ cally important carbonate cement dissolution, it is expected that such secondary porosity develops in partially rather than pervasively cemented sand-

13

stones. Mesogenetic waters probably do not migrate along a wide front, but are instead focused. Hence aggressive waters that cause cement dissolution in a sandstone unit are usually derived from deeper levels. Over the past 20 years a major debate has centred on the mesogenetic dissolution of carbonate ce­ ments. One side of the argument suggests that carbonate dissolution is caused by acidic waters and C0 derived on thermal maturation of organic 2 matter in mudstones (Schmidt & McDonald, I 979; Morton & Land, 1 987). On the other side, mass balance calculations suggest that the amounts of organic matter may be insufficient to provide nec­ essary C02 that could produce the observed carbon­ ate dissolution and secondary porosity seen in most sandstones (Lundegard & Land, 1 986). Moreover, acidic waters may be neutralized within the mud­ stones due to interactions with carbonate bioclasts and silicate minerals before reaching adjacent sand­ stones (Giles & Marshall, 1 9 86). Carbonate cement dissolution can also be accomplished by means of carboxylic acids and carboxylic acid anions formed by redox reactions during the hydrocarbon invasion of hematite-bearing sandstones (Surdam et a/., 1 993). Alternative mechanisms that account for the mesogenetic dissolution of carbonate cements in sandstones include: (i) the cooling of ascending hot waters aided by the retrograde solubility of carbon­ ates (Giles & de Boer, 1 990; Wood & Hewett, 1 9 84); and (ii) the mixing of two waters (Runnells, 1 969). The resulting saturation state of carbonate cements due to mixing depends on: Pco,, tempera­ ture, ionic strength (salinity), degree of carbonate saturation and pH of the end-member waters before mixing (Thraikill, 1 968; Plummer, 1 97 5 ; Wigley & Plummer, 1 976; James & Choquette, 1 990). Dissolution of carbonate cements in the shallow subsurface realm is attributed to the infiltration of meteoric waters, which are weak carbonic acids, or to mixing corrosion. The overall leaching capacity of meteoric waters is strongly controlled by: (i) the amounts of dissolved C02 available in the soil profile; (ii) the type and extent of organic-inorganic reactions that produce or consume protons; (iii) the permeability and depositional geometry of the sandstones; and (iv) the hydraulic heads. The disso­ lution capacity is expected to be more significant in permeable, laterally extensive sandstones in basins with a high hydraulic head. However, meteoric waters are unlikely to cause deep burial mineral

S. Morad

14

dissolution because they probably attain equilib­ rium with carbonates and silicates in the soil profile and in the relatively shallow subsurface. The more reactive the mineral contents in these regimes, the shallower is the dissolution capacity of meteoric waters. Morad et a!. (this volume) concluded that in the Triassic Lunde Formation, North Sea, meteoric waters dissolved carbonate cements and framework silicates within a few tens of metres below the Kimmerian unconformity surface. In areas where the Lunde Formation was buried at depths> 350 m below this surface, meteoric waters mainly caused the dissolution of carbonate cements. Criteria for the recognition of secondary porosity due to the dissolution of carbonate cements in sand­ stones include (Schmidt & McDonald, 1 979) the presence of: (i) oversized pores formed by the disso­ lution of grain-replacive carbonate cements-over­ sized pores may, however, result from the dissolu­ tion of carbonate bioclasts and intraclasts; (ii) partially dissolved carbonate cements with etched rather than euhedral crystal outlines; and (iii) grain­ replacive carbonates surrounded by open pores. Secondary cement dissolution porosity that mimics or enhances primary intergranular porosity is difficult to recognize. Nevertheless, dissolved car­ bonate cements may leave framework grains with corroded margins that can be best recognized by scanning electron microscopy (Burley & Kantorow­ icz, 1 986). Dissolution of calcite cement is more pervasive than the less soluble dolomite, ankerite and siderite.

RECRYSTALLIZATION AND REPLACEMENT OF CARBONATE CEMENTS

In addition to dissolution, the destabilization of car­ bonate cements may result in recrystallization and replacement by other carbonates. Microcrystalline calcite and dolomite are sensitive to recrystallization at various burial depths. The recrystallization of dolomite has been reviewed by Mazzullo ( 1 992). Burial recrystallization of micritic/microsparitic ce­ ments in sandstones may result in the formation of poikilotopic calcite (Saigal & Bj0rlykke, 1 9 8 7). However, poikilotopic calcite is also a common primary cement in calcretes (e.g. Knox, 1 977; Tan­ don & Narayan, 1 9 8 1 ). Recrystallized calcite and dolomite are recognized as patchily distributed, coarsened crystals. In contrast, precipitational vari-

ations in crystal size of drusy carbonates show trends of increasing crystal size from pore walls to pore centre. Siderite and ankerite are less soluble and thus less sensitive to recrystallization than calcite and dolomite (Matsumoto & Iijima, 1 98 1 ; Mozley & Bums, 1 992). Spot! & Bums ( 1 994) argued that magnesite is resistant to deep burial recrystallization, but might undergo recrystalliza­ tion by interaction with meteoric waters at low temperatures. Recrystallization may influence the crystal struc­ tural, elemental and isotopic compositions of the carbonate in question (Gregg et a!., 1 992; Chafetz & Rush, 1 994; Malone et a!., 1 994; Kupecz & Land, 1 994). Carbonate cements formed by recrystalliza­ tion are characterized by lower 8180 values than the microcrystalline precursor cements. This suggests the involvement of meteoric waters or increased burial temperatures. Therefore recrystallization must be considered when 8180 is used for studies on palaeoclimate, the timing of cementation and palaeo-water composition. Unlike 8180, the carbon and strontium isotopic compositions of carbonates may be preserved during recrystallization, particu­ larly in low permeability rocks (Dutton & Land, 1 98 5 ; Siegel et a!., 1 98 7 ; Cerling, 1 99 1 ; Driese & Mora, 1 993; Kupecz & Land, 1 994). In addition to recrystallization, the replacement of one carbonate cement by another is common during burial diagenesis. Eogenetic calcite cement may be replaced partially to completely by ferroan dolomite/ ankerite during mesodiagenesis (Boles, 1 978). The dolomitization of calcite cement, which is wide­ spread in limestones, is less frequently reported for sandstones (Hudson & Andrews, 1 9 87; Lawrence, 1 99 1 ; Morad et a!., 1 995). Complete replacement of calcite cement by dolomite and ankerite is difficult to recognize. However, its recognition may be possible by the presence of mimetically replaced bioclasts (Richter & Fuchtbauer, 1 9 78; Morad et a!., 1 996) and by the similarity of the dolomite and ankerite fabric to the eogenetic calcite. Replacement of sider­ ite cement and intraclasts by ankerite occurs in res­ ervoir sandstones from offshore Norway (Morad et a!., 1 996). Upon uplift and invasion by meteoric waters, dolomite/ankerite may be dissolved or re­ placed by calcite ± hematite (Morad et a!., 1 995). Calcitization of dolomite cements may also occur in the eogenetic regime due to subtle modifications in pore water chemistry caused by variations in inten­ sity of rainfall and/or sea- or lake-level fluctuations (e.g. Colson & Cojan, 1 996) (Fig. 4).

·

Geochemical evolution of carbonate cements EQUILIBRIUM RELATIONSHIPS AMONG DIAGENETIC CARBONATES

Studies on the stability of diagenetic minerals in relation to temperature and formation water chem­ istry provide important insights into the overall mineralogical and chemical evolution of the host sediments (Boles, 1 982; Kaiser, 1 984; Morad et a!., 1 990, 1 994). The precipation conditions and equilibrium relationships of carbonate are com­ plex issues and controlled by several inter-related parameters, such as pore water chemistry (ionic activities, pH, alkalinity, dissolved organic com­ pounds), kinetics and temperature. The tempera­ ture-dependent equilibrium relationships among calcite-ankerite-siderite, calcite-dolomite-mag­ nesite, siderite-magnesite, and dolomite-ankerite have been calculated as functions of aMgl+/Uca'+, 1.0) fall within the stability field of siderite in Fig. Sa. However, no siderite has been detected in these sediments. Low log (a.Fe, . fa.ca2 + ) ratios, and hence the plot of pore waters in the stability field of calcite, is often related to the precipitation of Fe-sulphides. The stability relationship and perhaps even the extent of solid solution between siderite and magne­ site depend on the temperature and the Fe/Mg activity ratio. As the temperature increases, the sta­ bility field of magnesite increases, which means that higher Fe/Mg activity ratios are required to stabilize siderite at the expense of magnesite (Fig. 5b). For­ mation waters from Triassic Tunisian reservoirs ( T � 80 " C; Morad et a/., 1994) are characterized by log (a.FeH/a.Mg, . ) ratios of -3.9 to -2.4 and hence fall within the stability field of siderite, which is far more dominant than magnesite (Morad et a/., 1994). Al­ ternating zones of magnesian siderite and ferroan magnesite in these Triassic sandstones formed at 5560 ' C (Morad et a/., 1994). Magnesium-rich siderites with low Ca and Mn contents have also been formed at increased temperatures (� 70-90 ' C) in other sedimentary basins (Macaulay et a/., 1993; Mozley & Hoernle, 1990; Rezaee & Rojahn-Schulz, this volume). When several generations of siderite occur in a sedimentary sequence, it appears that the later generations are more enriched in Mg (e.g. Mozley, 1989b). Iron-rich magnesite cements in Permian mudstones and sandstones from Austria have been reported by Spot! & Burns ( 1994). Eogenetic ferroan magnesite (FeC03 � 1.5-25.5 mol%) also forms in mudstones and sandstones of deep-sea sediments (Matsumoto & Matsuda, 1987; Matsumoto, 1992). Unlike magnesite formed at in­ creased temperatures (Morad et a/., 1994; Spot! & Burns, 1994), these deep-sea magnesites contain substantial amounts of Ca (6.5- 17.0 mol%) and Mn (0.5-20.5 mol%). Eogenetic magnesian siderites in marine sediments contain appreciable amounts of Ca (McKay et a/., 1995; Mozley, l 989a; see also Browne & Kingston, 1993; Morad et a/., this vol­ ume). Deep-sea siderites are enriched in Mn (Chow et a!., 1996). Conversely, near pure or slightly to moderately Mn-rich (�2- l0 mol%) siderites form during the eodiagenesis of continental sediments

(Mozley, l 989a; Browne & Kingston, 1993; Baker et a/., 1996). The stability relationship between dolomite and ankerite depends on the temperature and activity 2 ratio of Fe 2+/Mg + (Fig. 5c). Formation waters from Norwegian North Sea reservoirs (Egeberg & Aagaard, 1989) have log (a.FeH/a.Mg, . ) of � -3 to -2 and fall within the stability field of dolomite and ankerite (Fig. 5c). Both of these minerals are widely reported as mesogenetic cements in these sediments (Saigal & Bje�rlykke, 1987; Morad et a/., 1990). The stability relationships between calcite, dolo­ mite and magnesite depend on the temperature and activity ratio of Mg 2 + /Ca 2 + (Fig. 5d). Lower Mg/Ca activity ratios are required to induce the dolomitization of calcite and to stabilize magnesite at the expense of dolomite (Fig. 5d) (Usdowski, 1994). Formation waters from the Norwegian North Sea reservoirs have an average log (a.Mg2 + / Uc3H) � - 1.0 t o 0.0 and thus fall within the stability field of dolomite. Nevertheless, both calcite and dolomite are common cements in these rocks, indicating that dolomitization is a kinetically con­ trolled reaction. Further evidence of this is revealed from Recent sediments, such as the Fraser River delta in Canada (Simpson & Hutcheon, 1995) (log (a.MgH/Uc3H) � -2.2 to +1.0), where the pore wa­ ters are saturated with respect to dolomite, but it is calcite rather than dolomite that precipitates. Cal­ cite rather than dolomite forms below the deep-sea floor, yet the pore waters plot at shallow, near sea bottom temperatures in the stability field of dolo­ mite and shift with an increase in depth towards the stability field of calcite (Fig. 5d). This shift is due to a diffusion-controlled, downhole decrease in Mg/Ca activity ratio caused by the incorporation of Mg in Mg-silicate that results from the alteration of volca­ nic material, a process which is coupled with the release of calcium (McDuff & Gieskes, 1976).

PATTERNS OF FLUID FLO W : CLUES T O THE ORIGIN AND MECHANISMS OF MESOGENETIC CARBONATE CEMENTATION

There is ample evidence of active, large-scale fluid flow in the subsurface, which should be considered in diagenetic modelling (Sullivan et a/., 1990; Glu­ yas & Coleman, 1992; Gaupp et a/., 1993). Direct evidence of fluid flow is manifested by hot springs,

Geochemical evolution of carbonate cements geyser fields, seafloor vents and seepages, and a rise in groundwater level during and after earthquakes (Sibson, 1990). The most important regimes of fluid flow in sedimentary basins (Fig. 6) are related to compaction by sediment loading, tectonic compres­ sion, deep meteoric infiltration in areas of tectonic uplift, thermo-chemical convection due to density gradients around salt diapirs and convection due to the presence of thermal gradients, such as in the vicinity of rising magmas. Fracturing, folding and thrusting greatly influ-

Fig. 6. Patterns of fluid flow

envisaged for three common types of sedimentary basins.

17

ence the style and extent of fluid flow in sedimen­ tary basins. Faults, however, may either act as high permeability conduits and thus enhance fluid flow (Knipe, 1993) or as seals that result in compartmen­ talization, and thus the restriction of water flow (Harding & Tuminas, 1989; Hindle, 1989). Tec­ tonic stresses cause rapid, pulse-like changes in fluid flow (Sibson et a/., 1975; Muir Wood, 1993). Fluid flow along fracture systems is episodic and occurs by seismic pumping and seismic valving (Sibson, i 98 1 ). Seismic pumping occurs due to pressure

18

S. Morad

gradients, whereas flow by seismic valving occurs as a result of dilation and fault failure induced by high pore pressures in the vicinity of overpressured �ones. The release of overpressure may be accom­ panied by hydrofracturing and fluid migration along pressure gradients (Sullivan et a!., 1 990; Caritat & Baker, 1 992; Schulz-Rojahn, 1 993). Fracturing and fluid flow along pressure gradients may result in mesogenetic carbonate cementation in intergranular pores of sandstones and along frac­ tures according to any of the following mechanisms. I Decrease in Pco, induced when fluids migrate to high permeability, underpressured lithologies, such as at interface between mudstones and sandstones, or along fault zones that are connected to under­ pressured zones. The precipitation of calcite can thus be envisaged as follows: 2 Ca + + 2HC03- CaC03 + C02 + H20 =

In a manner similar to carbonate precipitation in fractures, wellbore-scale precipitation and forma­ tion damage occur due to pressure release in hydrocarbon-producing wells (Fisher & Boles, 1 987). 2 Addition of C02 may induce carbonate precipi­ tation when the pH is externally buffered. Migra­ tion of C02 occurs along pressure gradients either in gaseous form driven by buoyancy, or dissolved in water by diffusion or advection. C02 in sedimen­ tary basins forms by inorganic reactions and by organic matter maturation. Reservoirs containing large volumes of C02 may be formed by the metamorphism of calcareous sequences due to the emplacement of igneous intrusions (Studlick et a!., 1 990). C02 can also be produced as a consequence of the pervasive dissolution of carbonate cements and carbonate rocks (Lundegard & Land, 1 9 86). 3 Increase in HC03 concentrations due to the degradation of oil by incurred meteoric waters. This is evidenced by carbonate cementation along the oil-water surface, such as in Tertiary, turbiditic reservoir sandstones from northern North Sea (Watson et a!., 1 995). Although the presence of cements along fractures is indicative of water flow, precipitation does not necessarily occur by advection, but rather by ionic diffusion from the host sediments. Advective ce­ mentation requires the circulation of huge water volumes. For each pore volume of cement, 1 04 to 1 05 water volumes are required (Bathurst, 1 97 5 ; Wood, 1 986; Sharp e t a!., 1 988). The distinction between cements formed by diffusive and advective

material flux is difficult, but certainly important in mass transfer studies. In contrast with diffusion, advection may indicate the derivation of external waters that have been subjected to temporal varia­ tions in chemical composition. This would result in complex chemical zonations within the carbonate crystals. Mesogenetic carbonate cements are derived inter­ nally from within the sandstones and externally from interbedded and juxtaposed beds as well as from waters migrated from deeper parts of the basins along fractures. The dissolution and repre­ cipitation of eogenetic carbonate cements and bio­ clasts are among the important internal sources. Albitization of Ca-plagioclase has also been consid­ ered as an internal source of calcium (Schulz et al., 1 989), but probably accounts for a small portion of calcite cement in sandstone sequences (Morad et a!., 1 990). External sources include interbedded and tectonically lower or juxtaposed lithologies such as mudstones, carbonate rocks and evaporites (e.g. Purvis, 1 992; Gaupp et a!., 1 993). Evidence used in support of external sources includes a greater abundance of carbonate cements at the boundaries with adjacent mudstones (e.g. Carvalho et a!., 1 99 5 ; Moraes & Surdam, 1 993). However, Sullivan & McBride ( 1 99 1 ) found no relationship between carbonate cement distribution in sand­ stones and the mudstones of the Gulf Coast Ter­ tiary. Moreover, in the absence of pH buffering agents, waters charged with high Pco, derived from mudstones may indeed induce carbonate dissolu­ tion rather than precipitation. Ca-charged dolomi­ tizing waters derived from deeply buried carbonate rocks migrate upwards and contribute to the calcite cementation of sandstones (Morad et al., 1 994). Burial carbonate cementation occurs subsequent to considerable compaction, leading to a successive decrease of both intergranular volume(IGV) and of o 180 of the carbonate. However, in some basins, carbonate cementation may occur by ascending hot basinal brines to shallow depths (Sullivan et al., 1 990). Such cements occur in weakly compacted sediments and are characterized by low 8180 values and fluid inclusions with high homogenization tem­ peratures. This mechanism imposes difficulties in recognizing these cements from those formed by recrystallization at increased temperatures, as both mechanisms preserve a high, pre-cement porosity. A few workers (Giroir et a!., 1 989; Souza et a!., 1 995) argued that the early emplacement of calcite cement in sandstones of rift basins may take place

Geochemical evolution of carbonate cements from hot convected waters driven by the high geothermal gradients related to the oceanic open­ ing. The role of hot water circulation due to the emplacement of diabase on the fracturing and diagenesis of sandstones has been proposed by Girard et a!. ( 1 988).

DIRECTIONS FOR FUTURE RESEARCH

Although a considerable advance has been made in our understanding of clastic diagenesis and of car­ bonate cementation in particular, factors control­ ling the cementation of ancient shallow marine sandstones which lack present day analogues, such as eogenetic, strata-bound calcite-cemented, marine sandstones are unclear. What are the sources of calcite cement in these sandstones when carbonate bioclasts are totally absent? Has the global and regional change in ocean chemistry and pattern of circulation any impact on cementation of sand on the sea floor? The numerical modelling of patterns, extent and mechanism of water flow in the subsurface and their influence on the mineralogy, geochemistry and dis­ tribution of carbonate cements should be an area of further future research. Questions that need to be answered include the following. Are the basinal brines in equilibrium with the conductive thermal field of the basin, and are they static, actively flowing or moving only sluggishly? Are water move­ ments induced essentially by extrinsic tectonic and thermal factors? What are the sources of deep, mesogenetic cements? What are the importance and scales of advection versus diffusion in their forma­ tion? Are metamorphic, magmatic, and perhaps even mantle-derived waters, involved in the diage­ netic evolution of formation waters and host sedi­ ments? Can carbonate cement redistribution and its influence on the properties of deep reservoirs be quantified? Volcanic events have been frequent throughout most of geological history, yet there are few studies documenting their importance in sandstone diagen­ esis and carbonate cementation in particular. Is this related to difficulties in recognizing volcaniclastic sediments due to their rapid and extensive diage­ netic alteration? Future research in clastic diagenesis would ben­ efit from an interdisciplinary approach with respect to other water-related disciplines, such as igneous/

19

hydrothermal, metamorphic, structural, ore and hydrogeology. Our view of diagenetic evolution of sandstones is currently strongly biased towards the rapidly subsiding basins of the Gulf Coast of USA and the North Sea in north-west Europe. Studies should include a wider diversity of basinal settings to approach a more realistic picture of clastic diagenesis. Finally, the sharp line between scientists dealing with the diagenesis of mudstones and car­ bonate rocks should be removed, and instead we should learn from what have been achieved by them, for instance about coastal and nearshore pore water chemistry and diagenesis, and about the factors .and mechanisms of massive dolomitization of limestones.

ACKNOWLEDGEMENTS

I thank I.S. Al-Aasm, L.D. De Ros, W. Dickinson, Q. Fisher, C. Macaulay, J. Hendry and C. Spot! for constructive reviews of the manuscript. I am grate­ ful to the Swedish Natural Science Research Coun­ cil (NFR) for supporting my research activities.

REFERENCES

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Spec. Pubis int. Ass. Sediment. ( 1 998) 26, 27-5 1

Origin and spatial distribution of early vadose and phreatic calcite cements in the Zia Formation, Albuquerque Basin, New Mexico, USA J.R. B E C K N E R a n d P . S . M OZLEY

Department of Earth and Environmental Science, New Mexico Institute of Mining and Technology, Socorro, NM 87801, USA

A BSTRACT

The Miocene Zia Formation consists of sands and muds depos ted i in fluvial, aeolian and playa lake environments. Although much of the formation is poorly consolidated, resistant zones of calcite cementation are common. These range in size from isolated nodules to tabular cemented zones several metres thick t hat extend for over 2 km laterally .The calcite cemented zones are h ghly i complex, exhibiting a wide range of macroscop ci and m croscopic i textures and geometries .Af er t cons dering i a combination of microscopic, macroscopic and geochemical characteristics, we have inferred the environment of precipitation (i e.. pedogenic, vadose non-pedogenic, phreat ci ) of the pr ncipal i types of cementat on i Nodules . and rhizocretions with micrit ci fabrics and alveolar structures are inferred to be vadose carbonates. Ovoid or elongate concretions, characterized by blocky spar cements and preservation of primary sedimentary structures, are inferred to be phreatic carbonates .Most cemented units in the Zia Formation reflect characteristics of both phreatic and vadose zone cementation (e g.. 3 preservation of sedimentary structures plus rhizocret ons i and alveolar microtextures ). 81 C values for 18 vadose cement tend to be heavier and 8 0 values tend to be similar or slightly lighter than phreatic cements .813 C and 818 0 values for units with mixed features tend to have intermediate values. Most cementation types that exhibit a m xture i of features may reflect past fluctuations of the water table, where vadose cements were moved into the phreatic zone. V adose zone cementation occurred princ pally i in association with soil development, whereas phreatic zone cementation occurred preferent ally i in zones of high primary permeability .In many cases early vadose cements provided nucleation sites for later phreatic cementation .Tabular units in the Zia Formation are of en t laterally extensive, decreasing potential reservoir/a q u ifer q uality by forming significant barriers to vertical f uid l flow .These barriers could result in compartmentalizat on i of the reservoir/a q u ifer, and extensively reduce production if wells were screened on only one side of a cemented layer .

INTRODUCTION

involved and to determine the controls on the spatial distribution of diagenetic alterations. In this paper we examine controls on the origin and spatial distribution of early calcite cements in the Miocene Zia Formation of New Mexico, in which calcite­ cemented low-permeability zones can extend for several kilometres laterally. Unlike marine sediments, where early diagenesis typically occurs entirely within the phreatic (satu­ rated) zone, early diagenetic alterations in terres­ trial sediments occur in both vadose (unsaturated) and phreatic zones. Furthermore, in terrestrial

Understanding fluid flow in aquifers and hydrocar­ bon reservoirs requires an understanding of hetero­ geneities in porosity and permeability in the material. A number of workers have examined the influence of primary depositional controls on aqui­ fer heterogeneity (e.g. Weber, 1982; Anderson, 1989, 1990; Davis et a!., 1993). To date, however, few studies have examined the influence of dia­ genetic alterations on porosity and permeability heterogeneities. To predict the subsurface distribu­ tion of diagenetic alterations that influence flow, it is necessary to understand the diagenetic processes Carbonate Cementation in Sandstones Edited by Sadoon Morad © 1998 The International Association of Sedimentologists ISBN: 978-0-632-04777-2

27

J.R. Beckner and P.S. Mozley

28

sediments significant alterations can occur during pedogenesis. Thus, a fundamental problem in any study of early terrestrial diagenesis is identifying vadose versus phreatic alterations. Although there are many studies that have investigated pedogenic carbonate formation, few have focused on non­ pedogenic cementation. Fewer still have addressed the problem of differentiating among different types of cements. In the Zia Formation we have been able to infer the environments of cement formation and the relationships between these environments and the subsequent spatial distribution of cementation. Our principal conclusion is that cementation in the phreatic zone occurred preferentially in zones of high primary permeability, whereas vadose cemen­ tation occurred principally in association with soil development. Furthermore, pedogenic carbonates apparently served as nucleation sites for later phreatic cementation, leading to complex zones of mixed pedogenic and phreatic cements.

TERMINOLOGY

Because the terminology for early carbonate ce­ ments is complex and somewhat ambiguous, it is necessary to define the terms used in this study. Carbonate cements are subdivided into three prin­ cipal types: 1 Pedogenic carbonate is carbonate that precipi­ tated in an active soil (i.e. the precipitation was related to pedogenic processes such as weathering, evapotranspiration, biological activity, etc.). Most detailed studies of early (i.e. before significant burial diagenesis) calcite cementation in semi-arid and arid settings have been on pedogenic carbon­ ates (Gile et a!., 1 966; Reeves, 1976; Esteban & Klappa, 1 983; Klappa, 1983; Rabenhorst et a!., 1984; Machette, 1985; Monger et a!., 199 1 ; Mack et a!., 1993). 2 Vadose non-pedogenic carbonate is carbonate that precipitated in the vadose zone but is not related to pedogenesis. Vadose non-pedogenic carbonates have been reported in the literature (e.g. Carlisle, 1983; Goudie, 1 983; Semeniuk & Searle, 1 985; Wright & Tucker, 1 991 ) , although few criteria were described to distinguish them from pedogenic ce­ ments. 3 Phreatic carbonate is carbonate that precipitated by non-pedogenic processes in the phreatic zone. Terrestrial phreatic carbonates have been described by many workers (Mann & Horwitz, 1 979; Arakel &

McConchie, 1 982; Carlisle, 1 983; Netterberg, 1969; Arakel et a!., 1 989; Wright & Tucker, 199 1 ; Spot! & Wright, 1 992; Burns & Matter, 1995; Mozley & Davis, 1 996). The terms calcrete and caliche are frequently used to describe some of the cement types mentioned above. They are most often used to describe a variety of cryptocrystalline calcium carbonate de­ posits resulting from pedogenic processes, that eventually forms indurated masses (Gile et a!., 1966; Read, 1974; Reeves, 1976; Semeniuk & Meager, 198 1 ; Esteban & Klappa, 1983; Klappa, 1 9 83; Netterberg & Caiger, 1 983; Machette, 1985; Milnes, 1 992). The term calcrete has also been used to describe a wide variety of calcium carbonate deposits resulting from groundwater processes (Netterberg, 1 969; Mann & Horwitz, 1 979; Seme­ niuk & Meager 1 981; Arakel & McConchie, 1982; Carlisle, 1 983; Semeniuk & Searle, 1985; Jacobson et a!., 1 98 8; Arakel et a!., 1 989; Wright & Tucker, 1 99 1 ; Spot! & Wright, 1 992). These terms will be avoided, as they have been used in a variety of ways by different workers.

GEOLOGICAL SETTING

The Zia Formation is the basal rift filling unit of the Santa Fe Group in the Albuquerque basin, part of the 1000 km long Rio Grande rift of Colorado and New Mexico (Lozinsky, 1994). The 10-2 1 Ma Zia Formation is exposed in a 55 km long arc extending from the Rio Puerco in the west to 24 km north of Albuquerque, New Mexico (Gawne, 198 1 ) . The upper part of the Zia Formation ( 1 6- 1 0 Ma; Ted­ ford, 1982) was deposited during the period of most active rifting (Chapin & Cather, 1 994). The study site is on the western margin of the Albuquerque Basin, about 20 km from Albuquerque, on the King Ranch (Fig. I ). The Zia Formation in this area is typified by exposed, resistant, well cemented hori­ zons bounding poorly consolidated sediments. It can be divided into sand-dominated, aeolian (Piedra Parada) and fluvial-aeolian (Chamisa Mesa) mem­ bers; a mud-dominated, flu vial member (Canada Pillares Member); and the sand-dominated aeolian­ fluvial member (Unnamed Member; Gawne, 198 1 ; Tedford, 1982) (Figs 2 and 3). The lower contact of the Zia Formation is unconformable with the Eocene Galisteo Formation and the Crevasse Can­ yon Formation of the Cretaceous Mesaverde Group (Gawne, 1 98 1 ; Tedford, 1982). The upper contact

29

Calcite cements in the Zia Formation 1070

1060

360

New Mexico

0

35 0

Albuquerque Basin

Explanation

D Albuquerque Basi� '-- Basin boundary Fault-hachures on r'-downthrown side; dashed where inferred or buried.

� N Fig. 1 .

0

Map of the Albu q uer q ue Basin showing the King Ranch study area. Modified from Lozinsky ( 19 94).

is the Sand Hill fault, a major normal fault that offsets the Zia Formation and units of the Upper Santa Fe Group by about 600 m (Mozley & Good­ win, 1995a (Fig. 3)). Facies associations (Miall, 1990) were defined from a detailed analysis of lithofacies in the study area (Table I ; Fig. 4). The classification used for fluvial sediments is from Miall ( 1990) and Davis et a/. ( 199 3). The terms facies and facies/lithofacies association are also used to define aeolian sedi­ ments and sedimentary characteristics (Kocurek, 1981; Kocurek & Dott, 1 9 81; Porter, 1987; Chan, 1989). The symbols used for flu vial and aeolian facies associations (e.g. CH, OF, EC, ES in Table I and Fig. 4) were developed for this study. Palaeosol

20mi

1..__.. 1 "' '--11

0

20km

formation is a function of surface exposure time and landscape stability. Palaeosols are important to an understanding of depositional environments and ancient flood basin accretion rates (Leeder, 1975; Allen, 1 9 86; Atkinson, 1986; Kraus & Bown, 1986; Davies et al., 1993). Because of this they will be considered separately from crevasse splay and over­ bank deposits.

METHODS

Sections of the Zia Formation were measured along four transects to examine lateral and vertical varia­ tions in lithology and cementation (Fig. 3). Key

J.R. Beckner and P.S. Mozley

30

Lithology

Sample Locations

72895-6 72895-2 72895-3 72895-4 72895-5 72896-6a 72895-6b 72895-7 6295-4 81995-20 81995-19 81994-18 81994-17 81994-16 8594-16 72194-3 81994-15 81994-14 8594-14 8594-13 81994-9 81994-8 81994-7 81994-6 81994-5 81994-4 EXPLANT!ON 72195-2 81994-3 8594-12 81994-2 Muds 8594-11 81994-1 72195-1 � 122394-1 l:::::::l 122394-2 Silty Sand 122394-3 122394-4 r:.:7:l 122394-5 L::d 122394-6 Sand 122394-7 122394-8 � 122394-9 1395-1 � 122394-10 1395-2 Unconformable 1395-3 Contact . 1395-4 1395-5 Faulted Contact 1395-6 1395•7 1395-8 1395-9 1395-10 1395-11 8594-1 0 81895-1 r,=,=,.=,=,.=�=,=,-=,=;,=,.=,=,.=�=,=,-=,=,l\\\\' m�s�2 8594-8 8394-7 8594-7 81895-3 8594-6 8394-6 8594-5 8594-4A 8594-4 8394-5 8594-2 8394-4 8394-3 8594-1A

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Fig. 2. Generalized strat graphical i column of the Zia F ormation showing ages, lithologies, depositional en v ironments and locations of the samples. This stratigraphical column was constr u cted from the fo u r detailed col u m ns whose ozins ky ( 1 988) . locations are shown in Fig .3. Terminology and ages from Tedford ( 1 982) and L

beds were traced laterally throughout the study area to evaluate the continuity of cementation and vari­ ations in bedding thickness and cement morpholo­ gies. Cemented units were classified by outcrop morphology, surface textures and sedimentary structures. Seventy-six samples were collected along

the four measured sections for petrographical and geochemical analysis (Fig. 2). Laterally continuous units were sampled in more than one area to examine variations in petrographic and geochemi­ cal characteristics. Thin sections were made from most of the samples, which were impregnated with

31

Calcite cements in the Zia Formation 000000000000000000000000000000 000000000000000000000000000000

Parada Member

11;:11 liiil � �

�

Galisteo Formation Crevasse Canyon Formation Fault

COLUMN4

Fig. 3. Geological m ap of the study area showing locations of stratigraphical colu m ns. Modified fro mGawne ( 1 98 1 ) .

blue-dyed epoxy before thin section preparation to identify porosity. These thin sections were analysed for authigenic textures and mineralogy using a standard petrographic microscope under plain light, crossed-polarized light and cathodoluminescence. Mean grain size, sorting and roundness data were collected from outcrops and thin sections using visual comparators (grain size: Amstrat Inc.; sort­ ing: Pettijohn et al., 1987; roundness: Powers, 1953). The cathodoluminescence was performed on a microscope equipped with a MAAS/Nuclide model ELM-3 Luminoscope. A Chittick apparatus (modified from Dreimanis, 1962) was used to de­ termine the total percentage of calcite, and to test for the presence of other carbonates. The analytical precision based on I 0 samples is better than 3%. On selected samples a JEOL-733 Superprobe, equipped

with a high-resolution back-scattered electron de­ tector, X-ray mapping features and image analysis software, was used to determine elemental compo­ sition and zoning in cements. Sample operating conditions were 20 nA sample current and 1-10 Jlm beam diameter. Carbonate standards were used and sample totals are I 00 ± 2% for all values. Finally, a Finnigan MAT Delta E isotope ratio mass spec­ trometer was used to analyse carbon and oxygen isotope values for each sample. Carbon and oxygen values were measured from C02 gas liberated from whole rock samples using 100% phosphoric acid. Data are reported in parts per million (o/oo), relative to PDB for oxygen and carbon. The analytical preci­ sion, determined from six standards, is better than 0.1 o/oo for both carbon and oxygen.

Table 1.

w N

S u mmary o flithological in ormation f or f aci f se associations; t reminology modifi ed rom f Miall ( 1 990) and D avis et a/. ( 1 993)

Faci se association

Litho aci f se pr se n et

G o e m tery

Grain siz /esorting

C m e n e tation typ se

CH Chann le+ l v e ee

Trough cross-b d ed d e sand ( S t) planar laminat d e sand ( S p) low angl ecross b- d ed d e sand ( SI) horizontally laminat d e sand ( S h) rippl ecross-laminat d e sand ( S r) massiv esand ( S m) massiv ,e crud ley b d ed d e silts and muds (Fm ) fin ley laminat d e to rippl d e silts and muds (F I) laminat ed silt, sand, and clay (Fsc )

Tabular to l n e ticular 0.2-3 m thick 1 0 m to >2 km in lat real ex t n et

Fin eto coars ,e mod reat ley sort d e sand/sandston e

Typ el and typ e3 (phr aetic ) tabular units

Trough cross-b d ed d e sand ( S t) planar laminat d e sand ( S p) low angl ecross-b d ed d e sand ( S l) horizontally laminat ed sand ( S h) rippl ecross-laminat d e sand ( S r) massiv esand ( Sm )

Tabular, massiv ,e l n e ticular, to thin w d e g -eshap d e sands, 0.2-4 m thick 0.5 m to 0.5 km in lat real ex t n et

Poorly sort d e sands and silty sands

Thin sandston esh ee ts ar eusually w lel c m e n et d e

Massiv e, crud ley b d ed d e silts and muds (Fm ) fin ley laminat d e to rippl d e silts and muds (Fl ) laminat ed silt, sand, and clay (Fsc ) silts and clays w/rhizocr teions (Fr )

Tabular to thin and lobat e 0. 1 - 3 m thick 0.5 m to 0.2 km in lat real ex t n et

Muds and silts

Pala o e sols on sand (Ps ) pala o e sols on silts and clays (Psc ) silt and clays w/rhizocr teions (Fr ) massiv esand w/rhizocr teions ( Smr )

Tabular to discontinuous and patchy 0. 1 - l m thick 0. 1 t o > 1 km lat real ex t n et

Muds, silts, v ery fin eto m d e ium silty and clay ye sand/sandston se

EC Cross-stratifi d e a o e lian dun ebodi se

Trough cross-b d ed d e sand ( S t e) planar laminat d e sand ( S p e) low angl ecross-b d ed d e sand ( S l e) horizontally laminat d e sand ( S h e) rippl ecross-laminat d e sand ( Sr e) massiv esand ( Sm e)

Tabular, l enticular and w d e g -eshap d e 1 -3 m thick > 1 km lat real ex t n et

Fin eto low er coars ,e mod reat ley to w lel sort d e sand/sandston e

S catt re d e ovoid to leongat econcr teions and small typ e1 and typ e3 phr a etic tabular units

E S A o e lian sandsh ee t d p e osits

L ow angl ecross-b d ed d e sand ( S l e) horizontally laminat d e sand ( Sh e) rippl ecross-laminat d e sand ( S r e)

Tabular 1 -2 m thick > 1 km lat r eal ex t n et

Fin eto m d e ium mod erat ley to poorly sort d e sand/sandston e

Coars re lay res of te n orm f w lel c em n et d e typ e 1 and typ e3 (phr a etic )

ID I nt erdun ed p e osits

L ow angl ecross-b d ed d e sand ( S l e) horizontally laminat d e sand ( S h e) rippl ecross-laminat d e sand ( S r e) massiv esand ( SM )

Tabular, l n e ticular and discontinuous 0. 1 -0.5 m thick 1 0 m to > 1 km lat real ex t n et

All siz se, g n e really poorly sort d e or bimodal

S m all typ e1 and typ e3 (phr aetic ) tabular units

cs

Cr v e ass esplay d p e osits

OF Ov rebank fin se

p

Pala o esol horizons

�

(l:l

Poorly c m e n et d e, isolat d e nodul se, platy, and rod concr teions

Nodular, platy and rod concr teions, and typ e2 and typ e3 (vados e) tabular units

t:l:l

�

?;�

"" "' 1=:>

� l=:l.. :-tl Yl

� � N

Calcite cements in the Zia Formation

33

Sand Dominated Fluvial Environment

r:7':77l l:i:::2d

Crevasse Splay Deposits (CS)

D •

Overbank Fines (OF)

D

Cross-stratified Eolian Dune Deposits (EC)

Paleosol Horizon (P)

E3 Eolian Sheetsand � Deposits (ES) � Interdune � Deposits (ID)

Fig. 4.

S ch m e atic d p e ositional g o e m ter y o fth eZia Formation. Faci se ass ociations ar e rom f Tabl eI.

SANDSTONE PETROGRA PHY

Most of the Zia Formation in the King Ranch area can be classified as lithic arkoses (Fig. 5). The Zia Formation can be further subdivided into two distinct domains on the QFL diagram: one contains the lower Zia Formation (Piedra Parada, Chamisa Mesa and Canada Pillares Members; the other contains the Unnamed Member. The lower Zia Formation changes from a feldspathic litharenite

(Piedra Parada Member) to lithic arkoses (Chamisa Mesa, Canada Pillares Members). The Unnamed Member exhibits scattered compositions, but is differentiated from the lower members by greater amounts of feldspar (Fig. 5). Volcanic rock fragments of intermediate compo­ sition are generally the most abundant lithic frag­ ments, averaging 70-90% of all rock fragments (Fig. 5). Chert is the most common sedimentary rock fragment, although some units contain abun-

J.R. Beckner and P.S. Mozley

34 Q Su b a ko r se

Fig. 5. T renary plot o f

po istion by eco m s ton s asnd/ and mem b re o f th eZia F or m ation . S a m pl ethat plot sa sa lithar enite contain sa larg ea m ount o fd terital ol k f mF ification ro carbonat e. Cla ss ( 1 974 ) .

•

Unn a medM em b er IJ Ca n a d a Pilla resM em b er • Cha m i as M es a M em b er 6 Pi edr a Pa ar d a M ember

dant detrital carbonate (Fig. 5). These carbonate fragments resemble pedogenic carbonates and may be caused by erosion of the underlying pedogenic units. Most volcanic lithic fragments are fresh and well rounded; however, chemical alteration has removed unstable phenocrysts such as hornblende and plagio­ clase from some volcanic grains, leaving euhedral voids. More irregular voids indicate dissolution of aphanitic/glassy groundmasses. Potassium feldspars vary from fresh to deeply altered to clays. Most of the plagioclase is unaltered, and dissolution along cleav­ age planes is more common than alteration to calcite or clays.

TYPES OF CALCITE CEMENTATION

Calcite cementation in the Zia Formation is com­ plex, exhibiting a wide range of macroscopic and microscopic morphologies. Four principal types of isolated concretions, and three principal types of laterally extensive tabular units, were identified. A summary and description of facies associations, lithofacies types, lithologic data and cementation types is given in Table I . Descriptive data and interpretations for each cementation type are pro­ vided in Table 2. Details of spatial distribution and Iitholacies/Iithologic associations of these cementa­ tion types are shown in Figs 6 and 7.

Concretions

Nodules Nodules can be subdivided into two types. The first consists of small (0. 1 -5 em diameter) subspherical to irregular forms (Fig. 8A) and is common in reddened clays and clay-rich silty sands in overbank fine (OF), and palaeosol (P) horizons (Table I ; Fig. 6). Some of this first type of nodule exhibit two stages of concentric zonation, distinguished by a colour change from grey or greenish grey in the middle to pink on the outside. Dense micrite forms the usual matrix, and crystallaria (with some cir­ cumgranular forms) are common (Fig. 8B). The second type of nodule is roughly the same size and shape, but is characterized by oval grooved and tubular surface pitting (Fig. 8C). This type is more common in the silts and silty sands (crevasse splay (CS), overbank fines (OF), and palaeosols (P)) of the upper Unnamed Member. It has a micritic matrix, circumgranular cracks, micrite-spar, and some alve­ olar textures as well (Fig. 8D). A micrite-spar mi­ crotexture is where grains or groups of grains are coated with micritic cements, and the areas in between are filled with spar (16-50 �m diameter).

Ovoid and elongate concretions These concretions range from small ( 1 -4 em diam-

Table 2. Summ a ry of d secri p t iv ed a t an a d int re p r te a t ions of c m e n et a tion ty pe s in th eZ i aF orm a tion

Environm n et of p r cei p it taion

Cem n e t taion ty pe

Host lithology

Outcro pmor phology

Surf a c et x e tur se

Microt x etur se

Nodul a r concr teions

C l a ys , cl y a-rich silty s a nd

0. 1 -5 e m di a m te re ovoid to irr geul a r sh ape s

Smooth to p i tt d e , tub ed n a d groov d e

Micritic f b a ric , m n e iscus c em ents , circumgr a nul a r cr a cking , cryst lal rai a, a lv o el a rt x e tur se , gr a in dissolution

V a dos e

Ovoid to leong ta e concr teions

Fin eto co a rs es a nd

l-4 e m di a m te re ovoid to > l 0 m leong a t esh ape s

Smooth to w a rty

Poikiloto p ic to blocky s pa r

Phr ea tic

Pl a ty concr teions

C l y as , cl a y-rich silty s n ad

5-50 e m caross p l ta se th ta s eem to follow r elict t x e tur se

Smooth to p i tt d e , tub d e n a d groov d e

Micritic f a bric , m n e iscus c m e n e ts , circumgr a nul a r cr caking , lav o el a r t xetur es , gr a i n dissolution

V a dos e

� �

Rod concr teions

C l a ys , silty s and , n ad s a nd

0. 1 -5 em di a m te re , 3-50 e m long ; singl eor br a nching , th ni downw a rds

Mostly smooth , but som teim se p i tt d e , tub d e n a d groov d e

Micritic f a bric , circumgr a nul a r cr a cking , a lv o el a r t xetur se , gr a in dissolution

V d a os e

Ty pel t a bul a r c m e n et d e unit

F in eto co a rs es and

G n e re laly > l 0 m l ta re a l xet n e t , with p r se rev d e s d e im n e t ar y structur se ; sh a r plow re , n ad g n e re a lly sh a r pu ppe r bound rai se

Smooth to w a rty surf a c e

Poikiloto p i c to blocky s pa r

Phr ea t ic

'"' "'

�

"'

� � ;:;·

s. "'

N iS'

� ....

Ty pe2 t b a ul a r c m e n et d e unit

V rey fin eto m d e iumgr a in d e cl y a y e to silty s n ad

G n e re laly l 0-200 m l ta re la xet n e t ; m sasiv e, mottl d e, w a vy- p l tay , br ceci a t d e, t eepee, l a m in a r f ea t ur se ; sh a r pu ppe r a n d diffus e low er bound rai se

Smooth to p i tt d e , tub d e n a d groov d e

Micritic f a bric , m n e iscus c m e n e ts , circumgr a nul a r cr a cking , r a di a l s pa r, lav o el a r n ad f n e setr a l t xetur se , gr a i n dissolution

V a dos e

Ty pe3 ( p h r ea t ic ) t a bul a r c em n e t ed unit

Fin eto m d e ium-gr a in d e s a nd

G n e re laly > l 0 m l a t er la xet ent , with p r se rev d e s d e im n et a ry structur se p l us rod sh ape s

Smooth to w a rty , som teim se v rey irr geul ra with p i ts , tub se , a nd groov se

Blocky s pa r n a d s pa rry flo a ting gr a in t x e tur se

Phr ea tic >> v a dos e

Ty pe3 (v a dos )e t b a ul a r c em ent d e unit

V rey fin eto m d e iumgr a in d e cl a y ye to silty s a nd

M sasiv e, mottl nodul se , rods , som es d e im n et structur se p r se

Smooth to p i tt d e , tub ed a n d groov d e

Micritic f a bric , m n e iscus c em n e ts, circumgr a nul a r cr caking , lav o el a r n ad micrit -es pa r t xetur se

V a dos e>> p h r ea t ic

d e , with a nd p l a t se ; a ry rev d e

Q

� � (3 �

w U>

J.R. Beckner and P.S. Mozley

36

(/) w a: 1w

EXPLANATION

::;

LITHOLOGY

j< J I:J t::: � k::: 1 km) (Fig. 7). Calcite cementation textures are mainly blocky spar. Coalesced ovoid

to elongate concretions are commonly found on the tops of these units: they are usually less than 1 m in thickness and can extend for tens of metres laterally.

Type 2 (no sedimentary structures preserved) Type 2 tabular units lack original sedimentary structures and are often associated with reddened clays and clayey sands from overbank fine (OF), palaeosol (P) and interdune (ID) deposits (Table 1 ; Fig. 6). Micritic calcite i s the main cement, and micrite-spar textures, grain dissolution, alveolar structures, circumgranular cracking and meniscus cement are common. Type 2 tabular units are subdivided by outcrop morphology into massive, platy, wavy bedded, fractured and laminar types.

40

J.R. Beckner and P.S. Mozley

e b re. Notic eth emillim ter -esiz d e tub e( T ) a nd Fig. I 0. (A ) Pl tay concr teion from th emiddl eof th eUnn am de M m l e ra ein c n e tim ter se. (B) Photomicrogr ap h of a lv o el a r t xetur se (A ) from a groov e(G) structur se. Divisions on th esc a p l a ty concr teion in th eUnn a m d e M m e b re. (C ) Rod-sh ape d concr teions from n a ae oli n a s n a dston ein th eCh a m is a M se aM m e b re. Not eth ta s v e re a l of th erods br a nch a n d thin downw a rds. Divisions on th esc la e ra ein c n e tim ter se. (D) R a di a l s pa r (microcodium) microt x etur e.

The most common type 2 morphology is charac­ terized by massive bedding, with abundant branch­ ing or isolated rod structures and pitted, tubular and grooved surface textures (Fig. l l B). Lower contacts are usually gradational. This morphology is generally 0.3-1 m thick, and occasionally can be of great lateral extent (>I km) (Fig. 7). Some outcrops are thin ( l 0-20 em), platy or wavy bedded, with pitted, tubular and grooved surfaces (0.5-3 em diameter). These thin bedded units are generally less than l 0 m in lateral extent. Other outcrops are characterized by millimetre­ sized calcite-filled fractures that are in places irreg­ ular, unoriented and fenestral, and sometimes re­ semble small folds (Fig. l l C). Original sedimentary structures are generally not preserved. These units may also be associated with tubular, rod and platy concretions. These outcrops exhibit alveolar and

fenestral microtextures, and displacement laminae in thin section. Some outcrops have an irregular wavy laminar (3- 1 0 em thickness) morphology. Individual lami­ nae vary from l to 2 mm in thickness. Units usually have sharp upper and lower contacts. These forms exhibit abundant alveolar and fenestral microtex­ tures (Fig. 1 1 D).

Type 3 (tabular units with mixedfeatures) The above descriptions are pure end-member ce­ mentation types. However, most tabular cemented units in the Zia Formation show a mixture of characteristics of these end-members. Units that are closest in appearance to the type l end-member have excellent preservation of sedimentary struc­ tures, with rare pit and tube structures (type 3)

Calcite cements in the Zia Formation

41

Fig. 11. (A) Ty p e I t b aul a r unit from the middle of the Unn a med Member. Note the good p reserv taion of sediment ary

structures. Units on the sc a le a re in decimetres. (B) Ty p e2 t b a ul a r unit from the Piedr aP ra d a aMember. Note b a sence of sediment a ry structures. Units on the sc a l e rae in decimetres. (C ) Tee p ee structure from the middle of the Unn a med Member. Units on the sc a le a re in centimetres. (D) Photomicrogr ap h of fenestr a l/l a m in a r microtextures common in t a bul a r units with l a m in a r, brecci a ted a n d tee p ee outcro pmor p hologies.

(Fig. 12A). The most common, thickest and most laterally extensive units (>2 km) are those that are close in appearance to type I tabular units (Fig. 7). Mixed feature cements near the type 2 end-member are associated with more poorly sorted, finer grained layers and pit, tube and rod structures, with some evidence of the original sedimentary struc­ tures (type 3) (Fig. 128). Type 3 units also show a mixture of cement textures, including floating grain and micrite-spar types. Floating grain microtextures are usually char­ acterized by grains surrounded by drusy to iso­ pachous sparry cements, with the remaining void spaces filled with micrite or microspar (Fig. 1 2C). This type of cement is most commonly found in units near the type I end-member. The micrite-spar microtexture is most common in mixed units near the type 2 end-member (Fig. 12D). In these units the spar is generally equal to or more abundant than the micrite cements.

CATHODOLUMINESCENCE AND ELEMENTAL COM POSITION

Authigenic calcite varies from bright orange to non-luminescent, whereas detrital carbonate is a dull orange. Poikilotopic and blocky spar associated with ovoid and elongate concretions and type I tabular units is typically a dull orange to non­ luminescent. Some type I tabular units, and most type 3 (phreatic), show some zonation (bright orange to dull orange-red and non-luminescent). In most cases this is not visible under plane polarized light. Micritic cements are either a dull orange-red or non-luminescent. Spar-filled alveolar and fenes­ tral textures associated with these micrites are only luminescent along the edges. Oscillatory zoning (regular and irregular) in this spar occurs rarely. Although zonation is visible under cathodolumines­ cence, it is not visible using back-scattered electron imaging. Microprobe analysis shows that, regardless

42

J.R. Beckner and P.S. Mozley

Fig. 12. (A ) Ty pe3 ( ph r ea t ic ) t b a ul a r unit. Although th re eis good p r se rev taion of s d e i me nt a ry structur se , atub eth ta br a nch se downw a rds is shown by th e a rrow. Th eh amme r is app roxi ma t ley 1 8 emlong. (B) Ty pe3 (v a dos e) t b aul a r unit. Th e a rrows p o int to r leict s d e i me n t ray structur se. Divisions on th esc a l e ra ein d cei me t r se. (C ) S pa r- m icrit e microt x etur efro m aty pe3 ( p h r ea t ic ) unit. Fr ame work gr a i ns a r eco ta d e by dis pl caiv eiso pa chous s pa r, n a d th es pa c e b tew ee n is fill d e with micrit e. (D ) Micrit es pa r microt xetur efro m aty pe3 (v a dos e) unit . Gr a i ns n a d grou p s of gr a i ns ra eco ta d e with micrit ,e a nd th es pa c eb tew ee n is fill d e with s pa r.

of microtexture, the cements are very near the calcite end-member composition (Fig. 13). Magne­ sium is the main impurity, and even this is less than I mol o/o. Cements from the Sand Hill fault at the top of the section show slightly more magnesium than Zia Formation samples (Mozley & Goodwin, 1995a) (Fig. 13).

ISOTOPE GEOCHEMISTRY

The isotopic composition of the various calcite types does not vary greatly. Carbon isotope values (8 1 3 C) range from -3.0 to -5.5o/oo PDB, whereas oxygen isotope values (8 1 8 0) range from -7.3 to -13.6o/oo PDB (Fig. 14). 8 1 3 C values for nodular, platy, rod-shaped concretions and type 2 tabular units are generally heavier than those of other types

regardless of stratigraphical position (Fig. 14) There is also a weak upward stratigraphical trend of increased 8 1 3 C values in the Unnamed Member for type 1 and type 3 tabular units. 8 1 8 0 values for the lower part of the Zia Formation show no definite trend with stratigraphical position, but there is an increase in 8 1 8 0 values in type 1 and type 3 tabular units higher in the section within the Unnamed Member (Fig. 14). The highest value for Zia Forma­ tion cements (-7.3o/oo PDB) approaches the average value of the fault cements (-7.1 o/oo PDB) (Mozley & Goodwin, 1995b). Samples collected along a 500 m lateral traverse of a single cemented horizon that intersects the fault exhibit no consistent variation in 8 1 8 0 with distance from the fault. The sample closest to the fault (0.5 m) has the closest value to the fault cements (-7.3o/oo PDB). Type 2 tabular units and nodular, platy and rod-shaped concretions are .

Calcite cements in the Zia Formation

0 Zia cements (spar)

Fig. 13. Tern a ry di g ar a m showing com p osition of micrite, s pa r n ad S n a d Hill f u a lt cements from the study rae a . The sc a le of the p lot is a t 99 mol% Ca C 0 3 . D ta a for f u a lt cements from Mozley & Goodwin ( 1 99 Sb ).

generally more enriched in 1 3 C and depleted in 1 8 0 than those associated with type I and type 3 tubular units and ovoid and elongate concretions (Fig. 1 5).

D ISCUSSION Environments of cement formation

We have inferred the environments of cement formation in the Zia Formation by comparing microscopic and macroscopic characteristics with those of cements of known origin described in the literature. In this section we discuss known charac­ teristics of vadose and phreatic cements, and use this as the basis for identification of cementation environments in the Zia Formation.

Characteristics of vadose cementation Despite the complexities and variations in surficial environments of precipitation, vadose zone ce­ ments in arid environments have a number of distinctive characteristics. 1 A dense micritic fabric, crystallaria, circumgran­ ular cracking and alveolar textures have been fre­ quently associated with pedogenic cementation (Esteban & Klappa, 1 9 83; Wright, 1 990; Wright & Tucker, 1 99 1 ; Mora et a!., 1 993). Microcodium has also been associated with pedogenic cementation. Microcodium, which exhibits a radial spar micro-

43

texture, is associated with either root filaments, or casts of fruiting or resting stages of soil fungi (Klappa, 1 978, 1 979; Esteban & Klappa, 1 9 83; Goudie, 1 9 83; Wright, 1 990; Monger et a!., 1 99 1 ; Wright & Tucker, 1 99 1 ; Mora et a!., 1 993). 2 Permeability in the vadose zone tends to be higher in finer sediments because flow occurs preferentially along grain surfaces rather than the centre of large pores. Finer sediments have more surfaces on which vadose flow can occur (Palmquist & Johnson, 1 962; Hillel, 1 9 80; Jury et a!., 1 99 1 ; Mozley & Davis, 1 996). If cementation is limited by the supply of Ca 2 + and/or HC0 3- to the precipitation site, vadose cements should occur preferentially in the finer sediments (Mozley & Davis, 1 996). 3 Vadose cements are commonly associated with soil zonation and the alteration of parent material during soil development, resulting in reddened clays and clay-rich sands in which there is little or no preservation of original sedimentary structures (Retallack, 1 990; Mack et a!., 1 993; Mora et a!., 1 993). 4 Vadose cementation is intimately associated with rhizocretions, which record the orientation and position of former root systems as root casts or moulds (Klappa, 1 980b; Esteban & Klappa, 1 9 83; Goudie, 1 983; Retallack, 1 9 8 8 , 1 990; Wright & Tucker, 1 99 1 ; Gardner et a!., 1 992; Milnes, 1 992). 5 Vadose cementation is sometimes associated with distorted or disrupted bedding, such as brecciation and teepee structures. Brecciation can result from cracking and drying during dewatering, or cracking and dissolution when well indurated carbonate lay­ ers are disturbed by growing roots (Klappa, 1 9 80a; Esteban & Klappa, 1 9 83). Growing roots also play a role in the formation of some teepee structures, when expansion along a single layer forces sediment up­ wards (Klappa, 1 980a). Tepee structures can also arise from expansive calcite and/or evaporite min­ eral growth (Klappa, 1 980a; Warren, 1 9 82; Goudie, 1 983). 6 Cementation in the vadose zone can also result in irregular, wavy, laminar cement morphologies. Laminar cemented zones with abundant root traces and alveolar and fenestral microtextures are thought to result from root mats forming in the zone of capillary rise (Cohen, 1 98 2 ; Semeniuk & Searle, 1 985; Wright et a!. , 1 98 8). Laminar units high in the vadose zone may have etched upper surfaces due to exposure (Semeniuk & Meager, 1 98 1 ), or have fewer and more vertically oriented rhizocretions (Cohen, 1 982).

J.R. Beckner and P.S. Mozley

44 -2

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CARBON

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Fig. 8. (A) Photomicrograph of displaced 'floating' quartz grains in a sandstone cemented by calcrete with drusiform

rims followed by blocky pore-filling calcite (p); crossed polars; (B) photomicrograph of coarse blocky calcite engulfing and replacing kaolinite (k) within a large vug; crossed polars; (C) backscattered electron (BSE) image of a sandstone with poikilotopic calcite cement replacing pore-filling kaolinite (k), dissolved feldspars and bright siderite remnants; (D) BSE image of complex oscillatory precipitation of calcite and dolomite; rhombohedral dolomite (d l ) is replaced by calcite (c l ), which is covered by a laminated microcrystalline dolomite rim (d2) with some calcite intercalations (arrow), which is overgrown by ankerite (ak), followed by coarse pore-filling calcite (c2); (E) BSE image of a calcrete and dolocrete with microcrystalline calcite replacing dolomite (ca), covered in large pores by collomorphic laminated microcrystalline dolomite (d), followed by coarse pore-filling calcite (cb); (F) BSE image of a dolocrete with microcrystalline dolomite (medium grey) as displacive rims and expanding mica flakes, followed by coarse pore-filling calcite (bright).

Sheet-flood sandstones in the Snorre Field

Fig. 9 . Chemical composition o f diagenetic carbonates

from the Lower and Middle Members of Lunde Formation.

a few show irregular, weak outwards Fe increase. The o 1 80p08 values of cal cite range from -12. 4o/oo tO -4. 8o/oo, 0 1 3Cpos values from -5.6o/oo to + J . 6o/oo, and 87Sr/8 6 Sr ratios between 0. 7 1 1127 and 0 . 7 1 1655 (Table 1). Fluid inclusions are relatively rare in the poikilo­ topic cal cite and mainly single phased, which re­ mained so after freezing, indicating entrapment at temperatures :o;; s o o c (Goldstein & R eynol ds, 199 4). A total of eight two-phase inclusions with very small gas bubbles (i.e. high liquid/gas ratios) yielded a very narrow range of homogenization temperatures of 62-68 °C. These values were not corrected for pressure. The precise melting temper­ atures of the first and last ice crystals were not possible to obtain. The inclusions are rounded in shape, :o;;6 Jlm in diameter, and display no fluores­ cence under UV light. Dolomite and ankerite

Dolomite and ankerite together are second in abun­ dance (av. 8 volo/o) after calcite, and occur both as cement in sandstones and in dolocretes and cal­ cretes (up to ;;;. so volo/o). In the dolocretes, dolomite forms rims on detrital grains an d extensive inter­ granular cements composed of small euhedral to subhedral rhombs ( I 0 mol% FeC03 (Fig. 12a) (r2 -0. 76). Dolomites with < I 0 mol% FeC03 show no correlation between Ca and Fe (Fig. 12a) (r2 -0. 16). These features indi­ cate that in the crystal structure of ankerite and ferroan dolomite there is an increase in alternating =

=

=

=

68

S. Morad et al.

Fig. 10. (A) Scanning electron micrograph of euhedral rhombohedral dolomite crystals associated with finely crystalline kaolin in a sandstone; (B) BSE image of scattered dolomite rhombs (bright) in a sandstone with intergranular kaolin and kaolinization of feldspar (lower centre) and mica (centre, with expanded edges); (C) BSE image of dolomite rhombs (medium grey) engulfing and replacing kaolinized pseudomatrix (dark grey) and siderite crystals (bright); (D) scanning electron micrograph of small dolomite rhombs on top of clay-coated grains; (E) BSE image of zoned dolomite cement with external zones of ankerite composition (bright); (F) scanning electron micrograph of poorly shaped, flattened m icrocrystalline siderite. ·

69

Sheet-flood sandstones in the Snorre Field 50 D dolomite • Fe-dolomite/ankerite • siderite

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=

Siderite

Siderite is most abundant (up to ::::: 14%) in fine­ grained sandstones rich in mica and clay pseudo­ matrix. It occurs as small subhedral or flattened rhombs (< 3-15 J..L m) (Fig. l OF) that replaced the detrital clays, and expanded as well as replaced the mica flakes (Fig. l 3A). In coarser-grained sand-

stones, siderite also occurs as relatively large euhe­ dral rhombs (� 120 J..L m) around and within dissolved and kaolinized feldspars and within Fe-Ti oxides that have been dissolved and replaced by euhedral anatase (Fig. l 3B). Both the finely and coarsely crystalline siderites are partially dissolved, preferentially in the crystal cores (Fig. 1 3C). These intracrystalline dissolution pores are partially filled by authigenic chlorite. Siderite is engulfed by other carbonate cements (e.g. Fig. lOC), but in a few sandstones microcrystalline siderite occurs as rims around dolomite rhombs. The siderites are moderately to strongly enriched in magnesium (7 . 3-20.9 molo/o MgC03; av. 13. 7 %) (Table l ; Fig. l l ), and display no crystal-chemical zonation. The o 1 8 PDB values of siderite range from - l 6 . 9 o/oo to -8 . l o/oo, and the 0 1 3Cp08 values from -I 0. 1 o/oo to - l .2o/oo (Table I ).

Fig. 13. (A) BSE image of a biotite which has been expanded and partially replaced by microcrystalline siderite; (B) BSE image of coarse, partially dissolved siderite crystals (s) surrounding an anatase rim (at) after a dissolved detrital Fe-Ti mineral (partially filled by dolomite and siderite); the interstitial spaces are filled by feldspar- and pseudomatrix-replacing kaolinite (dark) and calcite (ca) with bright pyrite framboids; (C) BSE image showing a dissolved siderite crystal-the dissolution void is partially filled by chlorite; (D) BSE image of a sandstone with extensive kaolin replacement of feldspars, pseudomatrix and mica; there are partially dissolved feldspars and Ti-minerals (bright crown); (E) scanning electron micrograph of kaolinite vermicules made by thin, irregular-edged platelets which replaced clay pseudomatrix; (F) scanning electron micrograph of dickitized kaolinite vermicules-thick, euhedral dickite crystals grew between and replaced almost totally thin kaolinite platelets (arrows).

Sheet-flood sandstones in the Snorre Field Clay minerals

The diagenetic clay minerals are mainly kaolin and chlorite. Kaolin replaces feldspar, mica and pseudomatrix, and fills intergranular pores in the sandstones (Fig. 13D) as well as vugs, burrows and root moulds in the calcretes and dolocretes (Fig. 8B). SEM examination revealed that both kaolinite and dickite are present. Kaolinite occurs as thin, irregular-edged platelets that are stacked in vermicular aggregates with delicate texture, indicat­ ing an in situ authigenic origin (Fig. 13E). Kaolin­ ized micas display the typical expanded texture (Figs 1 OB and 13D); kao1inization is more intensive along terminations of the mica flakes. Compared with kaolinized feldspars and pore-filling kaolinite, kaolinite vermicules which have replaced pseudo­ matrix reveal less intercrystalline microporosity (Fig. 13E) and contain abundant microcrystalline remnants of precursor smectitic clays as well as iron and titanium oxides (Fig. 1OC). Dickite is distinguished from kaolinite by SEM and XRD (X-ray diffraction) analysis. Dickite oc­ curs as euhedral monoclinic blocky crystals and has typical XRD reflections at 4. 13 A, 3. 79 A, 2. 5 0 A and 2.33 A in randomly oriented samples. Dickite occurs together with pervasively etched remnants of kaolinite, from which it inherited the vermicular habit (Fig. 13F). Dickite crystals are much thicker (:::::; :;;. 1-8 Jlm) than kaolinite (< 1 Jlm), and show no etching. Morad et al. (1994) concluded that these textural features are indicative of kaolinite transfor­ mation into dickite via small-scale dissolution­ reprecipitation. Kaolinite is engulfed (and thus postdated) by calcite, dolomite and siderite, both in the sandstones and in the calcretes and dolocretes. Kaolinite totally engulfed by these carbonates is well preserved or only slightly dickitized. Dickite is covered by, and hence predates, authigenic chlorite. Both kaolinite and dickite are engulfed by quartz overgrowths and coarsely crystalline calcite. Chlorite replaces kaolinite, clay pseu domatrix, infiltrated clays, micas and heavy minerals. It oc­ curs as rims composed of platelets oriented perpen­ dicularly to grain surfaces. The rims were formed by replacing infiltrated smectitic clay coatings, which were originally oriented tangentially to grain sur­ faces (Fig. 14A) (see Moraes & De Ros, 1990). These infiltrated clays were presumably introduced into the vadose zone of alluvial continental sedi­ ments under semi-arid conditions by episodic floods (Walker et a/., 1978; Moraes & De Ros,

71

1990). Typically, mechanically infiltrated clays are originally detrital smectites formed under semi-arid weathering conditions (see Keller, 1970; Walker et a/., 1978). This is evidenced by the dominance of smectitic clays in the mudstone samples and in the mud intraclasts. Infiltrated coatings and derived chloritized rims are conspicuous, particularly in medium-grained sheet-flood sandstones (up to 2. 7 volo/o). Transformation of smectitic coatings into chlo­ rites has occurred through an intermediate stage honeycombed aggregates of mixed layers of chlorite/smectite (CIS). Chloritization is incom­ plete, leaving remnants of smectite and CIS clays beneath the chlorite platelets (Fig. 14A). Chlorite has also pseudomorphically replaced biotite fl akes and vermicular kaolinite aggregates, with preserva­ tion of the original stacked habit (Fig. 14B). Illite is a minor diagenetic constituent occurring as fibres closely associated with chloritized pseudo­ matrix and infiltrated coatings. The presence of honeycombed mixed-layer illite/smectite (liS) might indicate that illitization occurred via this intermediate stage. Quartz and feldspars

Quartz cement forms on average 2.6 %. , but is abundant (up to 9 . 7 %) in sandstones poor in detrital and authigenic clays. Quartz occurs both as over­ growths on detrital quartz and as prismatic out­ growths in the presence of relatively thick infiltrated clay coatings or authigenic clay rims (Fig. 14C). Quartz cements are covered by, but also cover and engulf, diagenetic carbonates, kaolin and chlorite (Fig. 14D), suggesting a recurrent precipitation dur­ ing burial diagenesis. Detrital plagioclase and K-feldspar grains are albitized, distinguished by the typical petrographic and chemical features characterized for the Upper Lunde Member by Morad et a/. ( 1990). Albitized detrital feldspars contain dissolution voids and are untwinned or show irregular blocky to tabular extinction. Detrital plagioclases are far less calcian (An < 10 mol%) than those analysed in the Upper Lunde (An � 28 mol%) by Morad et a/. ( 1990). The albitized K-feldspar grains are composed of a larger number of smaller lath-like albite crystals, arranged parallel to each other in two directions that presum­ ably reflect traces of cleavage planes. On average, diagenetic albite replaced 6.2 bulk rock-volume o/o, corresponding to :::::; 2/5 of the detri-

72

S. Morad et a!.

Fig. 14. (A) Scanning electron micrograph of chlorite platelets which grew perpendicularly on coatings of infiltrated

smectitic clays (background); (B) scanning electron micrograph of chlorite which replaced pseudomorphically and pervasively kaolinite vermicules; small remnants of corroded siderite crystals (s) which were probably involved in the reaction; (C) scanning electron micrograph of discontinuous quartz overgrowths and prismatic outgrowths on top of clay-coated grains; (D) scanning electron micrograph of late quartz outgrowths which engulf chlorite; (E) BSE image of an anatase rim (at) around dissolved Fe-Ti grain filled by zoned dolomite (dl); showing intergranular kaolin (dark), siderite (s) as well as finely crystalline and framboidal pyrite (py); (F) BSE image of coarse barite (white) engulfing and replacing kaolinite (dark) within a vugular pore rimmed by microcrystalline dolomite (dl).

Sheet-flood sandstones in the Snorre Field

tal feldspars which survived early dissolution an d­ kaolinization. Plagioclase grains were more affected by albitization (3. 1 relative to 2. 1 o/o remaining detrital plagioclase) than detrital K-feldspar (2.6 relative to 7 . 3% remaining K-feldspar). However, the greater abundance of K-feldspar may be due to preferential elimination of detrital plagioclase by other earlier diagenetic processes, such as dissolu­ tion, kaolinization and replacement by carbonates. Albite (Ab �100 mol%) also occurs as small (< 130 Jlm) discrete crystals associated with chloritic clays, and is engulfed by late quartz cements. The detrital K-feldspar grains show overgrowths (,;2. 3%) with ragged or sawtooth-like outline. In some cases the overgrowths occur around mouldic pores that resulted from the post-overgrowth disso­ lution of detrital K-feldspar cores. This is probably related to the near end-member composition of the overgrowths, which renders them more resistant to dissolution and albitization than the detrital core. Other diagenetic constituents

Hematite occurs sparsely in the fine-grained flood­ plain sediments, as tiny pigments that are either evenly distributed in the sediment or closely associ­ ated with infiltrated clay coatings around frame­ work grains, and as alteration products of detrital Fe-bearing minerals such as Fe-Ti oxides. Diagenetic Ti-oxides are more abundant than iron oxides in sandstones (up to 2.6 %). They occur as local aggregates of bipyramidal anatase crystals, apparently formed by the complete alteration of detrital Fe-Ti oxides, wh ich are commonly associ­ ated with siderite and ankerite (Figs 13B an d 1 4E) (see Morad, 1988) or are scattered in chloritized pseudomatrix and biotite. Pyrite averages 0.2 volo/o, and only in a few samples forms up to 1.3 volo/o. It shows two occur­ rence habits: (i) fine crystals (< 2 Jlm) or framboids scattered in kaolinized or chloritized detrital clays and micas, or engulfed by coarse carbonate cements (Fig. 15E); and (ii) coarsely crystalline (up to �200 Jlm across), intergranular replacive cement. Barite occurs as scarce, large crystals (up to 2 mm) filling vugs and cracks an d engulfing as well as replacing kaolinite and carbonate cements in dolocretes an d calcretes (Fig. 14F). In the sand­ stones, barite occurs as a few poikilotopic and small crystals which cover, and thus postdate, chlorite rims aroun d framework grains. Some epidote, monazite an d zircon grains show

73

reddish-brown envelopes of solid bitumen which were polymerized from oil by the radioactive emis­ sion of the grains. Their presence within present­ day water zones probably indicates that either the original oil column was thicker than at present, or that emplacement of oil was gradual, from the base to the top of the structure.

DISCUSSIO N

Paragenetic sequence and overall diagenetic evolution

The relative timing of the main diagenetic processes in Lower and Middle Lunde Members is presented schematically in Fig. 15. Because of the complex diagenetic patterns and burial histories (see Fig. 5), a precise timing cannot be achieved for all the diagenetic effects observed. N evertheless, the volu­ metrically important diagenetic processes occurred under early, n ear-surface conditions. Products of the first burial phase in the Jurassic, of the tela­ diagenesis during the Kimmerian uplift (Late Jurassic-Early Cretaceous) and of the second burial diagenesis phase (Middle Cretaceous to Recent) are volumetrically less significant than during eodiagen­ esis. Eodiagenesis and telodiagenesis

At near-surface conditions the eogenetic and teloge­ netic processes and products are strongly controlled by several interrelated parameters. These include the chemical composition of meteoric waters, cli­ mate, hydrological setting, rate of deposition versus erosion, as well as detrital composition, permeabil­ ity, biological activities and organic-matter content in the sediments and soil horizo11s. What follows is a discussion of the role of eodiagenesis and tela­ diagenesis on the overall diagenetic evolution of the Lower and Middle Lunde Members. Silicates

The climatic conditions and episodic flooding that characterized the depositional setting of the Lunde enhanced the infiltration of suspended clay particles and the formation of coatings around framework grains in the sandstones. Clay minerals formed by weathering processes in the hinterland under arid to semi-arid climatic conditions would be expected to

S. Morad et a!.

74 a p proxi m a t e time

(Ma)

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K-fe l d s p a r quartz compaction compaction

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be dominated by smectite, as were the infiltrated clays, clay pseudomatrix and mudstones. The near­ surface eogenetic interaction of meteoric waters with detrital minerals resulted in the formation of kaolinite at the expense of detrital feldspars and micas. Eogenetic kaolinite was conceivably formed during periods (either seasonal or several years' variability of climate) of increased rainfall, which were followed by dry conditions that enhanced the formation of calcretes and dolocretes. Evidence of kaolinite formation during eodiagenesis includes the engulfment of kaolinite and dissolved and kaolinized feldspar by calcretes and dolocretes. Our results indicate that kaolinite distribution in the Lunde Formation is not strictly controlled by the Kimmerian uplift and erosion. This is due partly to the formation of kaolinite during eodiagenesis and partly to the strong relationship between kaolinite abundance and detrital composition of the sand­ stones, particularly the original amounts of feldspars and mud intraclasts. Pervasive kaolinite formation, coupled with dissolution of calcite and dolomite ce­ ments, has been substantial in well 34/4- 1. In well 34/7 -A-3H sandstones, the top of which was buried deeper below the unconformity than that of well

, _ , _

Fig. 15. Simplified paragenetic sequence of the main diagenetic processes in the Lower and Middle Lunde sandstones.

34/4-1 (358 m and 24 m, respectively}, pervasive calcite and, to a lesser extent, dolomite dissolution and creation of secondary porosity was not accom­ panied by kaolinite formation. The meteoric waters were thus aggressive towards detrital silicates imme­ diately below the unconformity, but apparently re­ mained undersaturated only in relation to the car­ bonates at greater depths. Meteoric water incursion indicates the presence of considerable hydraulic head, as well as exposure of the Lunde above sea level. Grains and cement dissolution and kaolinization were conceivably enhanced by the humid cli matic conditions that prevailed in the area during the Late Jurassic to Early Cretaceous. In contrast to eoge­ netic kaolinite, telogenetic kaolinite replaces com­ pactionally deformed micas and clay pseudomatrix. Evidence indicating that this dissolution was tela­ genetic and not related to the second mesogenetic phase (Middle Cretaceous to Recent) (see Figs 5 and 15) includes the presence of later undissolved euhedral calcite and dolomite cements that post­ date q uartz overgrowths, chlorite rims and feldspar albitization. The presence of randomly scattered patches of carbonate cement left by telogenetic

Sheet-flood sandstones in the Snorre Field

dissolution promoted the preservation o f a loose sandstone framework and of telogenetic secondary porosity during the second burial phase. Inhibition of compaction in these sandstones is evidenced by the presence of undeformed ductile grains such as micas. Kaolinite is dominantly replacing feldspar, mud intraclasts and pseudomatrix. Replacement of the pseudomatrix indicates that kaolinite formation occurred, at least partially, by telodiagenesis during the Kimmerian uplift, subsequent to compaction caused by the first burial phase (J urassic) (Figs 5 and 1 5). Carbonates

The high intergranular volume (IGV) in carbonate­ cemented Lunde sediments indicates an early, pre-compactional timing. Siderite was among the first carbonates to precipitate, after feldspar disso­ lution and kaolinization and calcite and dolomite cementation. According to Mozley ( 1 989), the relatively high Mg content (av 1 3. 7 mol%) in the siderites should indicate precipitation from marine­ influenced pore waters. However, no marine­ related sedimentary facies were detected in the sequence, and it is believed that the sea might have been up to hundreds of kilometres away from the Lunde depositional sites during the Ladinian to Norian (Steel & Ryseth, 1 990; N ystuen & Fait, 1 99 5). Therefore, the elevated Mg content in siderite is considered primarily to reflect high aM8> + related to alteration of the detrital magnesian minerals, such as biotite, heavy minerals and smec­ titic mud intraclasts, by infiltrated meteoric pore waters. Indeed, siderite is associated with dissolved and kaolinized micas and clays, which perhaps indicates that even iron, as well as suitable pH values, were provided by these altered silicates (see Boles & Johnson, 1 984; Morad, 1 990). Such siderite is more enriched in Mg than that in the open pores, which supports our hypothesis. Cementation by Fe-poor calcite and dolomite occurred recurrently during eodiagenesis, as indi­ cated by the mutual partial replacement and by the presence of alternating rims of both minerals. Dis­ tinction between vadose and phreatic cementation is not easy (see Purvis & Wright, 1 99 1 ; Spot! & Wright, 1 992). However, the samples lack typical vadose features, such as meniscus and pendant cements, rhizocretions and glaebules (Esteban & Klappa, 1 983; Arakel & McConchie, 1 982). It is

75

therefore believed that cementation was accom­ plished in the phreatic zone. Additional evidence for this postulation is the coarse crystalline texture and the presence of crystal-chemical zonation in the carbonate cements. Moreover, the microcrystalline carbonate cements display a homogeneous lumines­ cence (see Plate 1 B,D) which reflects periodically homogeneous pore-water compositions more typi­ cal of the phreatic zone. Conversely, vadose cal­ cretes and dolocretes are expected to have patchy variations in luminescence as a result of periodic influx of waters into the sediments. The dominance of phreatic over vadose cementation may be due to extensive alluvial reworking and poor vegetation. Calcite cement in calcretes shows lower Mn and/or Fe contents than do the pre-compactional, poikilotopic calcite cements, indicating formation under generally more oxidizing conditions. How­ ever, the presence of small-scale CL zonations in eogenetic, vug-filling calcite is related to fluctua­ tions in aMn>+ in the pore waters, which probably took place in the sub-oxic phreatic zone. Microcrys­ talline dolomite in the dolocretes is characterized by dark red luminescence and low Mn and Fe contents, suggesting likewise more oxic conditions than those of the Fe-rich dolomites in the sand­ stones. The bright fl uorescence of microcrystalline calcite and dolomite cements in calcretes and dolo­ cretes is attributed to adsorbed organic matter from microbial remnants (see Dravis & Yurewicz, 1 98 5). The influence of microorganisms, such as bacteria, lichens and algae, in calcrete and dolocrete precip­ itation is indicated by the preservation of bacterial and algal cell remnants and calcified filaments in these deposits (see Phillips et a!., 1 98 7 ; J ones, 1 988; Folk, 1 99 3). The alternating bands of calcite and dolomite in the calcretes and dolocretes resemble those formed by mixing between marine and ]lleteoric waters (Ward & Halley, 1 98 5 ; Machel & Mountj oy, 1 986; Humphrey & Radjef, 1 99 1 ; Morad et a!., 1 992). However, there is no facies evidence of marine influence on the studied sequence, and diagenesis was thus fully meteoric. Therefore, the alternating bands are attributed to episodic fl uctuations in the amounts of rainfall and dilution of the pore waters, and shifting between dolomite and calcite equilib­ rium fields. Dolomite was formed during dry peri­ ods of increase in the Mg/Ca ratio of pore waters caused by water-sediment interaction (e.g. alter­ ation of biotite and mud intraclasts), coupled with evaporation. This is supported by the higher Sr

76

S. Morad et al.

contents in dolomite (up to :::; 7 00 ppm) compared with calcite (up to :::; 2 70 ppm). Similar ranges of 87Sr/8 6Sr ratios in dolomite and calcite, however, indicate a similar source of strontium. Watts ( 1 980) observed similar alternating bands of calcite and dolomite in pedogenic calcretes from the Kalahari Desert, which he attributed to mixing between fresh phreatic waters and more saline, vadose waters. Dramatic fl uctuations in the near-surface geochem­ ical environment due to climatic changes would explain the close succession and sometimes alterna­ tion of kaolinite, siderite, dolomite and calcite. The sources of eogenetic calcite and dolomite cements in alluvial sediments of the Lunde and similar successions elsewhere are often not immedi­ ately clear. This is particularly true when the strata are not associated with carbonatic bedrocks or bioclasts, and there is no evidence for the presence of extraformational carbonate rock fragments. When no such carbonate sources are visible, cement may be derived from rainwater, airborne carbonate dust, and from the breakdown of calcian silicates and Ca-bearing plants (e.g. Goudie, 1 983; R eeves, 1 976). In the U pper Lunde sandstones much of the detrital plagioclase, which is moderately calcian (An .;;; 2 8 mol%; Morad et al., 1 990), was dissolved and kaolinized during eodiagenesis, and was thus a likely source of calcium ions. Magnesium as well as calcium was also derived from the kaolinization of mud intraclasts. The kaolinization of detrital biotite was an additional source of magnesium as well as iron. Moreover, microcrystalline carbonate intrac­ lasts derived from the erosion and redeposition of palaeosol sections are common in the studied rocks. These intraclasts were probably important sources for the syncompactional to mesogenetic carbonate cements. Increases in ionic concentrations of groundwaters, and enhanced carbonate precipita­ tion, may have subsequently occurred by evapora­ tion under the overall semi-arid climatic conditions (see White et al., 1 963). Mesodiagenesis: role o f eogenetic minerals and temperature

The mesogenetic reactions in the studied rocks were largely controlled by increases in temperature and by the patterns of eogenetic and telogenetic modifi­ cation. Syncompactional to early mesogenetic mod­ ifications were accomplished during two burial phases (see Figs 5 and 1 5), yet assignment of at least some of the diagenetic events to a specific burial

phase might be difficult. Carbonates formed during these modifications include: (i) calcite precipitated as euhedral blocky crystals and overgrowths on eogenetic cements; and (ii) ferroan dolomite an d ankerite cements precipitated as thin zones around early non-ferroan dolomite and siderite, and as discrete, zoned blocky crystals in sandstones. The blocky calcite was affected by telogenetic dissolu­ tion, and was thus mainly formed during the first burial phase (Jurassic) (see Figs 5 and 1 5), whereas calcite overgrowths, Fe-dolomite and ankerite dis­ play no signs of dissolution, which indicates that they were formed during the second burial phase (Middle Cretaceous to R ecent) . The important mesogenetic silicates formed as a consequence of considerable increases in tempera­ ture during the second burial phase (Fig. 1 5) in­ clude dickite, albite, chlorite and q uartz. Dickite is formed almost exclusively by the replacement of eogenetic kaolinite, a process that occurs at :::; 801 30 " C (Ehrenberg et al. , 1 993; McAulay et al., 1 993; Morad et al. , 1 994). Albitization of plagio­ clase occurred apparently simultaneously with dic­ kite formation . This process can result in the forma­ tion of minor amounts of kaolin and calcite owing to the presence of calcium and excess aluminium in the detrital plagioclase, compared with authigenic albite (Morad et al., 1 990). We are, however, unable to distinguish precisely these particular calcite and kaolin byproducts from the abundant earlier-formed kaolinite an d calcite cements. Nevertheless, kaolin booklets are commonly closely associated with albi­ tized plagioclase, and some of the early calcite cements display minor overgrowths that might be formed by mesogenetic calcite addition as a consequence of plagioclase albitization. As the provenance of Lunde sediments has not changed considerably with time, the low amounts and lesser extent of anorthite solid solution of plagioclase compared with those of the U pper Lunde (Morad et al., 1 990) is attributed to a more pervasive albitiza­ tion and elimination of particularly the calcian plagioclases in the Lower an d Middle Lunde Members. Q uartz overgrowths are absent to minor in sand­ stones totally cemented by calcretes and dolocretes, but quite common in sandstones cemented by syncompactional carbonate cements, suggesting that part of the overgrowths formed during early compaction. Mesogenetic q uartz occurs as over­ growths and outgrowths that engulf earlier-formed minerals, including dickite, albite and chlorite.

77

Sheet-flood sandstones in the Snorre Field

Eogenetic feldspar dissolution and kaolinization are potential sources for the early-burial quartz over­ growths. Determination of the source for quartz outgrowths and overgrowths is beyond the scope of this study. The marked differences in diagenetic mineralogy between sediments of the two wells studied indicate variations in depositional facies and perhaps dif­ ferent diagenetic evolution pathways. In well 34/7 A-3H sandstones, chloritization occurs in the uppermost Middle Lunde Member and continues into the Upper Lunde sandstones. Chlorite covers, and hence postdates, albite and dickite. The smaller amounts of diagenetic kaolinite in well 34/7 -A-3H are attributed to less significant telogenetic dissolu­ tion of silicates because of the presence of Middle Lunde sandstones at a greater depth below the Kimmerian unconformity than the sandstones of well 34/4-1 (358 and 24 m, respectively). Although both cores display floodplain mud­ stones and sheet-flood sandstones, the Middle Lunde samples in 34/7 -A-3H are dominated by fine- to medium-grained sheet-flood sandstones. The elevated initial porosity and permeability of these sandstones compared with the mudstones, siltstones and very fine to fine-grained sandstones has perhaps allowed larger amounts of mechanically infiltrated clays, which are preserved as smectitic coatings and/or transformed into chloritic or CIS and liS rims in the sandstones. Porosity evolution: compaction versus cementation and reservoir implications

The overall large intergranular volume (IGV; av. 34.5%) and low packing values (average packing proximity index, Pp of Kahn, 1956; 25. 8%) indi­ cate that cementation occurred early and limited the compaction during subsequent burial. Near­ f surface cementation is evident iom grain displace­ ment, the presence of undeformed ductile grains such as micas within the cement, and the occur­ rence of intraclasts containing carbonate cements similar to those in the sandstones. The plot of IGV versus cement vol% in sandstones with IGV < 40% reveals that cementation was a much more impor­ tant agent of porosity destruction than compaction (Fig. 16). The low values of petrographic macroporosity (av. 1 1 . ( %), and petrophysical porosity (av. 17. 4%; range 0.05-28. 4%) and permeability (av. 42.6 mD; range < 0.01-672 mD) are due not only to compac=

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