Palaeoproterozoic magnesite: lithological and isotopic

Palaeoproterozoic magnesite: lithological and isotopic evidencefor playa/sabkha environments

VICTOR A. MELEZHIK*, ANTHONY E. FALLICK , PAVEL V. MEDVEDEVà andVLADIMIR V. MAKARIKHINà*Geological Survey of Norway, Leiv Eirikssons vei 39, 7491 Trondheim, Norway(E-mail: [email protected]) Scottish Universities Environmental Research Centre, East Kilbride, Glasgow G75 0QF, UKàInstitute of Geology of Karelian Scienti®c Centre, Russian Academy of Sciences, Pushkinskaya, 11,185610 Petrozavodsk, Russian Karelia, Russia

ABSTRACT

Magnesite forms a series of 1- to 15-m-thick beds within the »2á0 Ga

(Palaeoproterozoic) Tulomozerskaya Formation, NW Fennoscandian Shield,

Russia. Drillcore material together with natural exposures reveal that the 680-

m-thick formation is composed of a stromatolite±dolomite±`red bed' sequence

formed in a complex combination of shallow-marine and non-marine,

evaporitic environments. Dolomite-collapse breccia, stromatolitic and

micritic dolostones and sparry allochemical dolostones are the principal

rocks hosting the magnesite beds. All dolomite lithologies are marked by d13C

values from +7á1& to +11á6& (V-PDB) and d18O ranging from 17á4& to 26á3&(V-SMOW). Magnesite occurs in different forms: ®nely laminated micritic;

stromatolitic magnesite; and structureless micritic, crystalline and coarsely

crystalline magnesite. All varieties exhibit anomalously high d13C values

ranging from +9á0& to +11á6& and d18O values of 20á0±25á7&. Laminated and

structureless micritic magnesite forms as a secondary phase replacing dolomite

during early diagenesis, and replaced dolomite before the major phase of

burial. Crystalline and coarsely crystalline magnesite replacing micritic

magnesite formed late in the diagenetic/metamorphic history. Magnesite

apparently precipitated from sea water-derived brine, diluted by meteoric

¯uids. Magnesitization was accomplished under evaporitic conditions (sabkha

to playa lake environment) proposed to be similar to the Coorong or Lake

Walyungup coastal playa magnesite. Magnesite and host dolostones formed in

evaporative and partly restricted environments; consequently, extremely high

d13C values re¯ect a combined contribution from both global and local carbon

reservoirs. A 13C-rich global carbon reservoir (d13C at around +5&) is related to

the perturbation of the carbon cycle at 2á0 Ga, whereas the local enhancement

in 13C (up to +12&) is associated with evaporative and restricted environments

with high bioproductivity.

Keywords Carbon, dolomite, isotopes, magnesite, oxygen, Palaeoproterozoic,playa, red beds, sabkha, stromatolite.

INTRODUCTION

Despite the fact that magnesite is a rare mineral insedimentary rocks, it forms large-scale deposits,

of which only two major types are exploited atpresent: ultrama®c-hosted deposits of cryptocrys-talline magnesite (`Kraubath type') and deposits ofsparry magnesite within ancient marine platform

Sedimentology (2001) 48, 379±397

Ó 2001 International Association of Sedimentologists 379

carbonates (`Veitsch type'). Magnesite in lacus-trine sediments in the vicinity of ultrama®c rocks(`Bela Stena type') and metamorphosed ultrama®crocks with elevated magnesite content (`Greinertype') are of minor economic interest. In recentenvironments, magnesite occurs in coastal lakesand sabkhas and in continental playa lakes (Pohl,1989).

Sedimentary-hosted magnesite deposits, beingan exclusive feature of Neoproterozoic±Palaeo-zoic marine shelf sediments, may be considered,along with abundant associated dolomite, to be animportant component of Earth's evolutionaryhistory. Archaean (Schidlowski et al., 1975) andPalaeoproterozoic (Aharon, 1988) sedimentary-hosted magnesites have been described from veryfew localities. Magnesite mineralization fromPalaeoproterozoic sedimentary environments ofthe Fennoscandian Shield has not yet beenreported, although the possible presence of mag-nesite in Palaeoproterozoic dolostone sequenceshas long been recognized (e.g. Aksenov et al.,1975).

In this article, Palaeoproterozoic sedimentarymagnesite mineralization from the »2 billion-year-old Tulomozerskaya Formation (TF) of theNW Fennoscandian Shield (Fig. 1) is reported.

The magnesite mineralization occurs within theTF carbonates, which are marked by an extremeenrichment in 13C reaching +18& (Yudovichet al., 1991). The TF carbonate sequence is partof the 2á3±2á06 Ga Palaeoproterozoic positivecarbon isotope excursion (e.g. Baker & Fallick,1989; Karhu, 1993) and is marked by the greatestenrichment in 13C known from the Precambrian(Melezhik et al., 1999).

This paper addresses the following: (i) thenature of the magnesite mineralization; (ii) depo-sitional environments of the host carbonatesequence; and (iii) the possible signi®cance ofmagnesite for understanding the extreme 13C-en-richment of the host dolostones.

GEOLOGICAL BACKGROUND

The local geology, stratigraphy and lithology weredescribed in detail by Sokolov (1987). The TF isone of seven formations in the Palaeoproterozoicsuccession of the N. Onega Lake area. ThePalaeoproterozoic succession starts with basalpolymict conglomerates (Palozerskaya Formation;Fig. 1) unconformably overlying the Archaeansubstratum. The basal conglomerates are con-

Fig. 1. (a) Geographical and geological location of the study area (marked by a square). (b) Geological map (the samearea marked in the small box) of the northern Onega Lake area (simpli®ed from Akhmedov et al., 1993).

380 V. A. Melezhik et al.

Ó 2001 International Association of Sedimentologists, Sedimentology, 48, 379±397

formably overlain by the 50- to 120-m-thickJangozersakaya Formation, which consists of redterrigenous rocks characterized by cross-beddingand desiccation cracks. These rocks are conform-ably overlain by the 70-m-thick Medvezhegors-kaya Formation, which consists of ma®c lava withsubordinate pale grey and red, cross-beddedquartz arenites and ®ne-pebble conglomerates.The beds are conformably overlain by the TF,which is a 680-m-thick unit composed ofstromatolitic dolostones, red quartz arenites andsiltstones. This formation is, in turn, unconform-ably overlain by organic carbon-rich siltstonesand mudstones with subordinate dolostones ofthe 1500-m-thick Zaonezhskaya Formation. Thelatter is followed by the Suisarskaya Formation, a400-m-thick succession of basalts intercalatedwith numerous gabbro sills. A gabbro sill fromthe upper part of the Suisarskaya Formation has aSm±Nd mineral isochron age of 1980 � 27 Ma(Pukhtel' et al., 1992). The Palaeoproterozoicsuccession ends with the 190-m-thick Vashezers-kaya Formation comprising greywacke and arko-sic sandstones.

The entire sequence was deformed and under-went greenschist facies metamorphism during the1á8 Ga Svekofennian orogeny. The paragenesischlorite±actinolite±epidote re¯ects a temperatureof 300±350 °C (Sokolov, 1987).

LITHOSTRATIGRAPHIC POSITIONOF MAGNESITE BEDS

The magnesites cannot be distinguished from thehost dolostones in hand samples and have beenfound as a result of systematic analysis. Five bedsof magnesite and magnesite-bearing dolostoneshave been detected (Fig. 2). Four layers are indrillhole 5177 at depths of 553á5, 568á5, 598á7 and799á0 m. In drillhole 4699, magnesite-bearing do-lostones occur at depths of 537á5 m and 758á0 m(Fig. 2). Most of the beds are apparently <2 mthick, whereas the major magnesite bed (5177-553á5) is approximately 15 m thick. The lateralextent of magnesite-bearing beds is unknown. Thethickest magnesite bed is sandwiched between twobeds of dissolution-collapse breccia (Fig. 2).

PETROGRAPHIC DESCRIPTIONOF MAGNESITE

The magnesite and magnesite-rich dolostone arewhite, grey or yellow. The magnesite beds are

mainly composed of structureless micritic (5177-568á5, Fig. 3), crystalline (5177-553á5, Fig. 4) andcoarsely crystalline varieties. Finely laminatedmicritic, stromatolitic magnesite is rare (5177-598á7).

The structureless micritic magnesite consists ofmicritic magnesite (5±10 lm) with remnants ofdolomite. Rare idiomorphic crystals (0á1±0á3 mm)of magnesite and dolomite developed in themicritic aggregate (Fig. 3) indicate recrystalliza-tion related to neomorphic/metamorphic proces-ses. Micritic magnesite is also found in sparryallochemical dolostones with syntaxial dolomitecement. Here, small magnesite crystals are presentwithin the dolomitic intraclasts and are absent inthe syntaxial dolomite cement and crystallinedolomite matrix. If clear syntaxial overgrowthsformed mainly in a burial environment (e.g.Tucker & Wright, 1990), then the magnesitecrystals probably replaced dolomite before mostof the burial carbonate cements formed.

Micritic magnesite replaced dolomite pseudo-morphically, producing very ®nely intergrownmagnesite±dolomite aggregate (Fig. 4). From themicritic dolomite±magnesite relationship, it isobvious that magnesite could only have originatedduring early diagenesis if dolomite is a primaryprecipitate. Alternatively, magnesites may haveformed somewhat later if dolomite is an earlydiagenetic replacement of a calcite precursor.

The structureless crystalline magnesite consistsof 0á1±0á5 mm idiomorphic or subidiomorphicmagnesite crystals apparently replacing eithermicritic aggregate of 15±10 lm magnesite crystals(Fig. 4) and subordinate dolomite rhombs or siltydolomitic micrite. The coarsely crystalline mag-nesite has a larger grain size (0á5±2á0 mm) butexhibits similar relationships to the earlier phasesof micritic dolomite and magnesite.

Micritic stromatolitic magnesite has undergonerecrystallization. When ®ne lamination is pre-served, the thin, pale grey laminae are predom-inantly composed of micrite, whereas thicker,light laminae consist of microsparite.

LITHOFACIES AND DEPOSITIONALENVIRONMENTS

Ten lithofacies are recognized in the TF sequence(Table 1, Fig. 2). These lithofacies and theirdepositional environments have been describedby Melezhik et al. (20001 ), who demonstrated thatthe TF was deposited in a variety of environ-ments. Terrigenous `red beds' present throughout

Palaeoproterozoic playa/sabkha magnesite 381


the sequence formed in three main depositionalsettings: (i) a braided ¯uvial system over a low-energy, river-dominated coastal plain (lithofaciesI); (ii) a low-energy, barred, evaporitic lagoon(lithofacies II and IV±VII); and (iii) a non-marine,playa lake (lithofacies III and VIII). Biostromaland biohermal columnar stromatolitic dolostones(lithofacies X) are abundant and formed inshallow-water, low-energy, intertidal zones,barred evaporitic lagoons and peritidal evaporiticenvironments. The red, ¯at-laminated stromato-lites formed in evaporative ephemeral ponds,

coastal sabkhas and playa lakes. The presence oftepees, mudcracks, halite casts, pseudomorphsafter calcium sulphate and abundant `red beds' inthe sequence suggests that terrestrial environ-ments dominated over aqueous, with partial ortotal decoupling between the stromatolite-dom-inated depositional systems and nearby sea.

The magnesite has been designated as litho-facies IX (Melezhik et al., 20002 ).

Further details will only be given for thoselithofacies that are closely associated with the ®vemagnesite beds documented.

Fig. 2. Lithofacies and position of magnesite and magnesite-bearing dolostone beds in the TF intersected by drill-holes 5177 and 4699. Magnesite beds are marked by grey bars.



Bed 5 (Fig. 2) is the major magnesite occurrence.It is sandwiched between the dissolution-collapsebreccia of lithofacies VIII. The magnesite±collapsebreccia assemblage is overlain by lithofacies III,which consists of intercalated brown, mud-cracked, haematite-rich siltstones and pink, platy,cross-strati®ed, dolomite-cemented quartz are-nites. Fine laminae in lithofacies III typically formlow-relief hummocks and swells with amplitudesof less than 1 cm and wavelengths of 1±3 cm.A 1-to 2-m-thick bed of clastic haematite ore andhalite cube casts in siltstones have also beenreported (Akhmedov et al., 1993). Lithofacies III issimilar to the terrestrial `red bed'±dolostone±halite association in the Bitter Spring Formationin the Amadeus Basin, Australia, which formed ina series of shallow, hypersaline lakes and ponds(Southgate, 1986).

The presence of halite casts and the absence ofcalcium sulphate evaporites is also consistentwith a non-marine origin for lithofacies III. Theabundant desiccation cracks and the presence ofsmall-wavelength wave ripples suggest depos-ition in a playa lake.

The dissolution-collapse breccia, which hoststhe Bed 5 magnesite, appears as poorly cementedfragments of brown, pink and white dolostones

and brown, ®nely laminated mudstones eitherembedded in insoluble residues or cemented bycoarsely crystalline dolomite (Fig. 5A). Theinsoluble residues are composed of dark browndolomite±sericite±chlorite material enriched iniron oxide. As no palaeokarst surfaces have beenobserved, the subsurface dissolution of evaporiteminerals was probably the main process that ledto the development of the collapse breccias.Sedimentological data match a playa lake orsabkha environment (Melezhik et al., 20003 ). Theclose spatial relationship of Bed 5 to lithofacies IIIand VIII indicates that it formed either in a playalake or in a ponded tidal ¯at setting underevaporitic conditions.

Bed 3 rests on lithofacies V and is overlain bylithofacies X. Bed 4 is entirely associated withlithofacies V, which consists of variegated, struc-tureless or indistinctly parallel-laminated dolo-stones, marls and mudstones. The rocks aresporadically marked by desiccation cracks andby dolomite pseudomorphs after displacive,isolated, small crystals of gypsum (Akhmedovet al., 1993). Some of these pseudomorphs have`swallow tail' twin morphology (Fig. 5B). Thedesiccation cracks are ®lled with quartz sand, areseveral decimetres wide and penetrate 2±3 cm.

Fig. 3. Photomicrographs showinglarge neomorphic crystals of mag-nesite (marked by `M') and dolomite(marked by `D') and micritic mag-nesite replacing micritic dolomite.(A) Plane-polarized light; (B)back-scattered electron image;(C) Ca X-ray map; (D) Mg X-ray map.Drillhole 5177, depth 568á5 m. 1 and2 in (B) are positions of electronmicroprobe analyses shown inTable 36 .



Syndepositional deformation is expressed astepee structures (Akhmedov et al., 1993). Thelack of post-depositional compaction of the des-

iccation cracks suggests that the lithi®cation oc-curred before burial (Melezhik et al., 2000). Thisis consistent with the depositional environment,

Fig. 4. Photomicrographs showing large crystals of magnesite replacing aggregate of micritic magnesite and dolo-mite. (A) Plane-polarized light; (B) back-scattered electron image; (C) Ca X-ray map; (D) Mg X-ray map. (E±G)Magni®ed views of intergrown magnesite±dolomite mixture. (E) Back-scattered electron image; (F) Ca X-ray map; (G)Mg X-ray map. Drillhole 5177, depth 553á5 m. 1±3 in (B) and 4±8 in (E) are positions of electron microprobe analysesshown in Table 37 .



Table

1.

Lit

hofa

cie

san

dth

eir

pala

eoen

vir

on

men

tal

inte

rpre

tati

on

.

Lit

ho-

facie

sR

ock

ass

em

bla

ge

Rock

colo

ur

Bed

din

gan

dla

min

ati

on

Cem

en

tS

peci®

cst

ructu

ral

featu

reS

uggest

ed

pala

eoen

vir

on

-m

en

tal

inte

rpre

tati

on

XM

icri

tic

an

dsp

arr

yst

rom

ato

liti

cd

olo

ston

es

Red

,beig

e,

wh

ite

Lam

inate

dN

ot

pre

serv

ed

Tep

ee,

tep

ee-r

ela

ted

bre

ccia

s,d

esi

ccati

on

cra

cks

See

text

IXM

icri

tic,

cry

stall

ine

an

dst

rom

ato

liti

cm

agn

esi

te

Wh

ite,

pale

yell

ow

Mass

ive,

lam

inate

dN

ot

pre

serv

ed

±P

laya,

sabkh

aen

vir

on

men

t

VII

ID

olo

mit

e-c

oll

ap

sebre

ccia

sB

row

n,

red

±S

eri

cit

e,

ch

lori

te,

haem

ati

te±

Su

bsu

rface

dis

solu

tion

of

hali

te/s

ulp

hate

inp

laya,

sabkh

aen

vir

on

men

tV

IIM

icri

tic

all

och

em

ical

dolo

ston

es

Red

Th

inly

lam

inate

dN

ot

pre

serv

ed

All

och

em

sare

exclu

sively

dolo

mit

icooli

tes

Low

-en

erg

yti

dal

zon

eof

pro

tecte

dla

goon

VI

Sp

arr

yan

dm

icri

tic

all

och

em

ical

dolo

ston

es

Red

Str

uctu

rele

ss,

cru

dely

stra

ti®

ed

Syn

taxia

ld

olo

mit

esp

ar

Gyp

sum

pse

ud

om

orp

hed

by

cau

li¯

ow

er-

like

qu

art

zaggre

gate

s

Low

-en

erg

y,

evap

ori

tic,

pro

tecte

dla

goon

VC

ryst

all

ine

dolo

ston

es

Beig

e,

pin

kS

tru

ctu

rele

ss,

ind

isti

nct

para

llel-

lam

inate

dN

ot

pre

serv

ed

Desi

ccati

on

cra

cks,

dolo

mit

e-

pse

ud

om

orp

hed

gyp

sum

,te

pee

Up

per

tid

al

zon

eof

evap

ori

tic,

pro

tecte

dla

goon

IVD

olo

mit

e-r

ich

san

dst

on

es

an

dsi

ltst

on

es

Beig

e,

pale

pin

kH

err

ingbon

ecro

ss-

bed

ded

,¯

ase

r-bed

ded

Dolo

mit

e±

Low

-en

erg

yti

dal

an

dsu

pra

tid

al

san

d¯

at

III

Haem

ati

tesi

ltst

on

es

an

dsa

nd

ston

es

Bro

wn

,re

dP

laty

,cro

ss-s

trati

®ed

Dolo

mit

eM

ud

cra

cks,

low

-reli

ef

hu

mm

ocks,

small

wave

rip

ple

s,h

ali

tecast

s

Pla

ya

or

pon

ded

tid

al

¯at

un

der

evap

ori

tic

con

dit

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s

IIS

an

dst

on

es

an

dsi

ltst

on

es

Gre

yT

hin

-bed

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or

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ticu

lar-

bed

ded

Dolo

mit

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qu

art

z±

Occasi

on

all

y¯

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ed

sup

rati

dal

zon

eon

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dal

¯at

IQ

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zare

nit

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y,

red

Rip

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-mark

ed

,cro

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or

stru

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Qu

art

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seri

cit

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han

nels

,ta

bu

lar

sets

of

cro

ss-s

trati

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on

,asy

mm

etr

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pp

les

Bra

ided

¯u

via

lsy

stem

on

acarb

on

ate

coast

al

pla

in



which was affected by long-term subaerial expo-sure, allowing the carbonate sediments to becomelithi®ed. The observed `swallow tail' twin mor-phology of dolomite pseudomorphed gypsumcrystals closely resembles those reported fromancient rocks elsewhere (Rubin & Friedman,1977; Spencer & Lowenstein, 1990). The litho-facies V gypsum appears to have precipitatedpenecontemporaneously with the dolomite fromshallow, near-surface brines (Melezhik et al.,2000). Because intrasediment gypsum growth pro-vides unequivocal evidence of post-depositionalcrystallization in an evaporitic environment(Demicco & Hardie, 1994), similar conditions areproposed for lithofacies V. Overall, the sedimen-tological data suggest that the micritic dolostonesof lithofacies V and the associated magnesites ofBeds 3 and 4 formed in a shallow-water, evapo-rative setting. The tepee structures suggest thatthe carbonate rocks were subaerially exposed forextended periods.

Bed 2 magnesite has close association withlithofacies I, II and VI. Lithofacies II does notcontain reliable genetic information. However,the main diagnostic features of lithofacies I, suchas the dominance of sand, lack of silty and muddyparticles, unidirectional cross-strati®cation, gen-erally ®ning-upward sequence and presence oferosional channels, are consistent with braided¯uvial systems over a low-energy, river-domin-ated coastal, carbonate plain. Lithofacies VI is ared grainstone, which is typically structureless or

crudely strati®ed by variations in colour andgrain size. Allochems include unsorted, roundedand angular intraclasts of dolostone and sporadichaematite and siliceous oolites. Abundant vugs,voids and cauli¯ower-like aggregates of quartz(with both castellated margins and mammillatedsurfaces) with crude radial fabric are common.Given the lack of micrite in lithofacies VI grain-stones, the depositional setting must have hadsuf®cient current or wave energy to winnow awaythe ®ne matrix (Folk, 1962). However, theunsorted, mixed rounded and angular intraclastssuggest that the depositional environment hasonly occasionally been in¯uenced by such cur-rents and waves. Another diagnostic feature isabundant cauli¯ower-like aggregates of quartz,which are pseudomorphs after calcium sulphatenodules (Melezhik et al., 2000). The discovery ofsolitary anhydrite nodules, tens of millimetres to0á25 m in diameter, in the Holocene sediments ofthe Persian Gulf (Curtis et al., 1963; Shearman,1966) led to the development of `carbonate±evaporite' or `sabkha' depositional models toexplain some ancient shallow-marine carbonatedeposits (Shinn, 1983; James, 1984; Hardie &Shinn, 1986). Thus, overall, the data suggest thatlithofacies II and VI, and the associated magnesiteBed 2, formed in a barred, river-dominated,coastal, evaporative carbonate plain, occasionallyin¯uenced by tidal currents and waves.

Bed 1 magnesite is hosted by lithofacies X red,haematite-rich, ¯at-laminated stromatolites that

Fig. 5. (A) Dissolution-collapsebreccia consisting of fragments ofbrown, pink, massive and ®nelylaminated mudstones cemented bycoarsely crystalline dolomite.Lithofacies VIII, drillhole 5177,depth 542á0 m. Core diameter is42 mm. (B) Dolomite-pseudomor-phed crystals of gypsum exhibiting`swallow tail' twin morphology.Lithofacies V, drillhole 5177, depth570á0 m. Core diameter is 42 mm.



form relief sheets. The stromatolite sheets arevery often cracked and characterized by eitherindistinct, clotted or ribbon fabric with polygonalprism cracks. The presence of blisters, clottedfabrics with fenestrae and abundant desiccationcracks suggests that the ¯at-laminated stromato-lites and associated magnesite Bed 1 formed indrained depressions and ephemeral ponds on acarbonate ¯at (playa lake or sabkha environment).

SAMPLES AND ANALYTICAL METHODS

Magnesite and dolomite samples were obtainedfrom drillcores. The drillholes 5177 (35°25¢00¢¢E,62°14¢29¢¢N) and 4699 (35°28¢00¢¢E, 62°14¢30¢¢N)were made by the Karelian Geological Expedition.Both drillholes are 800 m deep; they partlyoverlap and intersect the entire thickness(680 m) of the TF.

Whole-rock oxygen and carbon isotope analy-ses were carried out at the Scottish UniversitiesEnvironmental Research Centre using the phos-phoric acid method described by McCrea (1950)and modi®ed by Rosenbaum & Sheppard (1986)for operation at 100 °C. Carbon and oxygenisotope ratios were measured on a VG SIRA 10mass spectrometer. Calibration to internationalreference material was through NBS 19, andprecision (1 r) for both isotope ratios is betterthan � 0á2&. Oxygen isotope data were correctedusing the fractionation factor 1á00913 for dolo-mites and 1á00933 for magnesite recommended byRosenbaum & Sheppard (1986). The d13C data arereported in per mil (&) relative to V-PDB and thed18O data in & relative to V-SMOW.

Because the magnesite and dolomite are com-monly intergrown, the chemistry (XRF) and stableisotopic composition of the samples wereobtained by whole-rock analysis, with additionalelectron microprobe (EMP) measurements todetect small-scale chemical variation. A sequen-tial acid reaction was used in an attempt toresolve isotopic composition of different minera-logical components in ®nely intergrown mixturesof magnesite and dolomite. The approach wasbased on the procedures recommended by Al-Aasm et al. (1990). A three-step sequential disso-lution with phosphoric acid was used on 10 mgaliquots: (i) 2 h at 25 °C to react calcite; (ii) 24 h at50 °C to react dolomite; (iii) 2 days at 100 °C toreact magnesites.

The major and trace elements were analysed byX-ray ¯uorescence spectrometry at the GeologicalSurvey of Norway using a Philips PW 1480 X-ray

spectrometer. The accuracy (1 r) is typicallybetter than 2% of the oxide present (SiO2,Al2O3, MgO, CaO), even at the level of 0á05wt%, and the precision is almost invariablyhigher than the accuracy. The analytical uncer-tainties (1 r) for Sr, MnO and Fe2O3 are betterthan � 5á5 p.p.m., �0á003% and �0á01% respect-ively.

Back-scattered electron imaging, EMP measure-ments and cathodoluminescence of carbonateminerals were carried out at the Institute of theContinental Shelf in Trondheim. A Jeol 733 SEMMicroprobe (Noran Instruments) with a silicon/lithium detector and Norwar window type wereused. The operating conditions were as follows:take-off angle 40°, acceleration voltage 15 kV,beam current 15 nA, beam diameter 1 lm, work-ing distance 11 mm. A Proza-type matrix correc-tion (fourth generation) was used. The detectionlimit for Ca, Mg, Fe and Mn is 0á1 wt% (energydispersive spectrometer). The following calibra-tion standard references were used: calcite for Ca,dolomite for Mg, magnetite for Fe, bustamitefor Mn.

GEOCHEMICAL RESULTS

Fifteen whole-rock samples of dolostone and foursamples of magnesite from drillholes 5177 and4699 were analysed for oxygen and carbonisotopes as well as for major and trace elements(Table 2). Results of electron microprobe analysiscarried out on the selected samples are presentedin Table 3. Additionally, a sequential acid reac-tion with subsequent oxygen and carbon isotopeanalyses was used for 28 composite dolomite±magnesite samples (Table 4).

Dolostone

Mg/Ca ratios of the dolostone dolomite, based onEMP measurements, range from 0á46 to 0á63(Table 3) with an average of 0á55, which is lowerthan that for stoichiometric dolomite (0á62).

The spread in oxygen and carbon isotopevalues is from 17á4& to 26á3& (mean21á5 � 1á7&) and from +7á1& to +11á6& (mean+9á6 � 1á3&) respectively. Cross-plots reveal nostatistically signi®cant covariation between d13Cand d18O (Fig. 6). There is no relationshipbetween the Mg/Ca ratio and C/O isotopes, norbetween Mn, Mn/Sr and C/O isotopic values. Srand d13C and Sr and d18O show positive covari-ation, r � 0á70 and r � 0á48 respectively.



Table

2.

Wh

ole

-rock

ele

men

tal

an

dis

oto

pic

com

posi

tion

of

dolo

mit

ean

dass

ocia

ted

magn

esi

tefr

om

the

TF

.

Dep

th(m

/Sam

ple

)L

ith

olo

gy

SiO

2

(%)

Al 2

O3

(%)

Fe

2O

3

(%)

MgO

(%)

CaO

(%)

Na 2

O(%

)M

nO

(%)

F (%)

Sr

(p.p

.m.)

d13C

(&)

d18O

(&)

Mg/

Ca

Ca/

Sr

Mn

/S

r

Dri

llh

ole

5177,

dolo

ston

es

542á0

Dis

solu

tion

-coll

ap

sebre

ccia

1á3

±0á0

221á9

31á8

0á1

40á2

50á4

3219

9á0

18á6

0á5

81030

8á7

9542á5

Dis

solu

tion

-coll

ap

sebre

ccia

2á8

±0á0

221á5

30á9

0á1

40á1

70á4

1185

8á9

19á6

0á5

91186

7á0

8554á5

Dis

solu

tion

-coll

ap

sebre

ccia

±±

0á0

222á8

32á4

0á1

30á0

80á4

4169

7á9

21á8

0á5

91361

3á6

4555á5

Cry

stall

ine

dolo

ston

e1á5

±0á0

222á5

31á4

0á1

30á0

70á4

5158

8á1

21á8

0á6

11410

3á4

1556á0

Cry

stall

ine

dolo

ston

e0á2

±0á0

222á4

32á2

0á1

30á0

70á4

6166

8á0

22á2

0á5

91375

3á2

5560á5

Cry

stall

ine

dolo

ston

e7á5

±0á0

219á3

28á3

0á1

40á0

70á4

0118

9á3

21á7

0á5

81702

4á5

7563á0

Str

om

ato

liti

cd

olo

ston

e13á2

±0á0

217á6

25á2

0á1

30á0

10á3

483

9á3

17á5

0á5

92152

0á9

3584á5

Sp

arr

yall

och

em

ical

dolo

ston

e34á1

0á2

50á1

414á1

20á5

0á1

30á0

90á2

582

7á9

21á1

0á5

81778

8á4

5593á0

Str

om

ato

liti

cd

olo

ston

e7á6

±0á0

319á6

28.3

0á1

40á0

50á3

8181

8á2

21á7

0á5

91109

2á1

3599á0

Cry

stall

ine

dolo

ston

e3á9

0á4

50á2

422á0

29á5

0á1

30á0

60á4

3193

8á3

22á5

0á6

31085

2á3

9788á0

Cry

stall

ine

dolo

ston

e8á9

0á6

10á2

419á0

27á4

0á1

6±

0á5

767

10á6

22á8

0á6

82904

±796á5

Cry

stall

ine

dolo

ston

e8á7

±0á0

619á7

27á4

0á1

30á0

50á3

7318

10á7

22á0

0á5

9612

1á2

1

Dri

llh

ole

5177,

magn

esi

te553á5

Cry

stall

ine

magn

esi

te14á0

±±

26á1

0á8

±0á1

7±

59á0

23á1

26á5

1179

261á8

0568á5

Mic

riti

cm

agn

esi

te±

±±

30á6

3á5

±0á1

9±

11

9á4

24á3

7á3

72266

133á0

0598á7

Str

om

ato

liti

cm

agn

esi

te8á2

1á6

00á3

526á8

9á1

±0á0

5±

62

9á6

21á4

2á5

01039

6á2

1799á0

Coars

ely

cry

stall

ine

magn

esi

te2á3

1á2

00á4

331á5

8á6

±0á0

6±

71

11á3

20á9

2á6

3860

±

Dri

llh

ole

4699,

dolo

ston

es

522á5

Sp

arr

yall

och

em

ical

dolo

ston

e15á3

0á1

00á0

817á1

24á8

0á1

40á0

40á3

255

9á6

21á6

0á5

83195

5á6

0524á1

Sp

arr

yall

och

em

ical

dolo

ston

e7á8

±0á0

219á4

28á3

0á1

50á0

60á4

375

9á3

20á9

0á5

82681

6á1

6527á5

Sp

arr

yall

och

em

ical

dolo

ston

e4á5

±0á0

621á4

29á6

0á1

30á0

40á5

5491

10á8

23á0

0á6

1429

0á6

3530á0

Sp

arr

yall

och

em

ical

dolo

ston

e1á5

±0á0

422á0

31á4

0á1

30á0

40á5

3449

10á9

23á3

0á5

9496

0á6

9535á5

Sp

arr

yall

och

em

ical

dolo

ston

e14á7

±0á0

217á5

24á6

0á1

20á0

40á3

4396

11á1

22á0

0á6

0440

0á7

8

Dri

llh

ole

4699,

magn

esi

te-b

eari

ng

dolo

ston

e537á5

Cry

stall

ine

magn

esi

te1á5

±0á0

232á8

10á6

±0á0

5±

151

11á2

21á0

2á6

2497

2á5

5

±,

Belo

wd

ete

cti

on

lim

it.

Dete

cti

on

lim

itis

<0á1

for

SiO

2,

<5

for

Sr

an

d<

0á0

1fo

roth

er

ele

men

ts.



Table

3.

Ele

ctr

on

mic

rop

robe

an

aly

ses

(wt%

)of

carb

on

ate

min

era

lsfr

om

the

Tu

lom

ozers

kaya

Form

ati

on

.

4699-5

06á0

5177±553á5

Dolo

mit

eall

och

em

Cem

en

t,sy

nta

xia

ld

olo

mit

eC

ryst

all

ine

magn

esi

teM

icri

tic

dolo

mit

e

Com

pon

en

tA

vera

ge

Avera

ge

1w

2w

3w

Avera

ge

4w

5w

6w

Avera

ge

MgO

*37á9

338á2

238á0

836á5

937á4

637á0

397á9

697á1

798á6

797á9

336á1

035á7

334á2

535á3

6F

eO

*0á3

40á0

00á1

70á1

90á0

00á1

00á0

00á5

70á0

00á1

90á1

40á1

90á2

40á1

9C

aO

*60á7

361á5

561á1

462á4

861á7

262á1

00á2

00á4

30á2

20á2

860á8

562á5

563á0

962á1

6M

nO

*0á0

00á0

00á0

00á3

20á3

10á3

20á7

30á8

30á7

20á7

60á6

90á9

10á8

00á8

0S

um

99á0

099á7

799á5

899á4

998á8

999á0

099á6

197á7

899á3

898á3

8M

gC

O3

42á3

042á2

842á2

940á6

841á6

241á1

599á2

698á5

499á2

599á0

240á8

839á8

838á6

939á8

1F

eC

O3

0á2

90á0

00á1

50á1

60á0

00á0

80á0

00á4

40á0

00á1

50á1

20á1

60á2

10á1

6C

aC

O3

57á4

157á7

257á5

658á8

858á1

258á5

00á1

70á3

70á1

90á2

458á4

059á1

860á4

159á3

3M

nC

O3

0á0

00á0

00á0

00á2

70á2

70á2

70á5

70á6

50á5

60á5

90á6

00á7

80á7

00á6

9S

um

100á0

0100á0

0100á0

0100á0

0100á0

0100á0

0100á0

0100á0

0100á0

0100á0

0M

g/C

a0á6

30á6

30á6

30á5

90á6

10á6

0493á2

0227á6

0451á6

0340

0á6

00á5

80á5

50á5

7

Table

3.

Con

tin

ued

.

5177-5

53á5

5177-5

68á5

5177-5

84á5

5177-7

96á5

Mic

riti

cm

agn

esi

teD

olo

mit

eC

ryst

all

ine

magn

esi

teZ

on

ed

eu

hed

ral

dolo

mit

eM

icri

tic

magn

esi

te

Com

pon

en

t7

w8

wA

vera

ge

1w

2w

Overg

row

thE

dge

Core

Avera

ge

MgO

*96á9

997á7

997á3

937á5

798á8

232á3

235á7

333á9

798á4

398á1

498á2

9F

eO

*0á1

60á0

00á0

80á0

00á0

00á0

00á0

00á0

00á3

60á0

00á1

8C

aO

*0á7

70á4

50á6

161á6

00á1

566á5

962á3

663á9

40á9

20á9

10á9

2M

nO

*0á5

30á0

20á2

80á0

90á3

90á0

00á9

41á1

20á0

40á4

10á2

3S

um

98á4

598á2

699á2

699á3

698á9

199á0

399á0

399á7

599á4

6M

gC

O3

98á7

999á6

099á1

941á8

199á5

736á4

140á0

138á1

698á9

198á9

098á9

1F

eC

O3

0á1

20á0

00á0

60á0

00á0

00á0

00á0

00á0

00á2

80á0

00á1

4C

aC

O3

0á6

60á3

90á5

358á1

10á1

363á5

959á1

860á8

70á7

80á7

80á7

8M

nC

O3

0á4

20á0

20á2

20á0

80á3

00á0

00á8

10á9

70á0

30á3

20á1

7S

um

100á0

0100á0

0100á0

0100á0

0100á0

0100á0

0100á0

0100á0

0100á0

0M

g/C

a126á9

0218á8

0173

0á6

1663á4

00á4

90á5

80á5

4107á7

0108á6

0108



Magnesite

The whole-rock XRF analyses indicate that the®ve beds of magnesite and magnesite-bearingdolostones are characterized by Mg/Ca ratios thatincrease upwards through the sequence from 2á6to 26á5 (Table 2). This indicates that the concen-tration of magnesite increases upwards throughthe sequence with the culmination in uppermagnesite Bed 5. The magnesites have lower Naand F contents than the dolomites (below detec-tion limits; Table 2). They are also depleted in Sr(<5±151 p.p.m. averaging 57 p.p.m.) comparedwith the dolomites. However, in the case ofmagnesite, Sr having a larger ionic radius shouldsubstitute only for Ca (similar to dolomite; Kretz,1982; Reeder, 1983). If the Sr concentrations arenormalized to Ca, there is no difference betweenmagnesites and host dolostones.

EMP measurements suggest that the Mg/Caratios of the magnesite correlate with the petrog-raphy of the rocks. Large magnesite crystals(neomorphic/metamorphic magnesite) have highMg/Ca ratios (230±663), whereas the Mg/Ca ratiosof micritic magnesite drop to 108±172, whichcorresponds to 0á78±0á53 CaCO3 wt% (Table 3).The overall CaCO3 content of the magnesite islow, in the range 0á13±0á78 wt% (Table 3). Themicritic magnesite is marked by a relativelyenhanced CaCO3 content (0á53 wt% on average)compared with the crystalline and coarsely crys-talline neomorphic magnesite (0á24 wt% onaverage). The crystalline and coarsely crystallinemagnesite is relatively enriched in Mn comparedwith micritic magnesite.

d13C and d18O values range from +9á0& to+11á6& (mean +10á1 � 1á0&) and from 20á0& to25á7& (mean 22á2 � 1á7&) respectively. Overall,the magnesites and their host dolostones arerather similar in carbon isotope ratios (Fig. 6,Table 4). However, the Bed 4 magnesite is char-acterized by slightly elevated d18O values (25á0&on average) compared with the average hostdolostone (21á3&, Table 2) and intergrown dolo-mite (21á6&, Table 4).

EVALUATION OF DIAGENESISAND METAMORPHISM

There is growing evidence that dolomite in thePrecambrian precipitated either directly from seawater (Tucker, 1982) or by dolomitization duringearly diagenesis caused by waters isotopicallycomparable with sea water (e.g. Veizer & Hoefs,T

able

3.

Con

tin

ued

.

5177-7

96á5

Core

,m

icri

tic

dolo

mit

eR

ecry

stall

ized

coate

dd

olo

mit

egra

inM

an

tle,

mic

riti

cd

olo

mit

eR

im,

mic

riti

cd

olo

mit

eC

em

en

t,sy

nta

xia

ld

olo

mit

eM

atr

ix,

dolo

mit

esp

ar

Com

pon

en

tA

vera

ge

Avera

ge

Avera

ge

MgO

*30á8

731á9

231á4

035á8

735á1

436á1

135á7

132á7

031á8

032á2

538á6

538á4

9F

eO

*0á0

40á1

50á1

00á3

70á0

00á0

00á1

20á0

00á0

00á0

00á0

00á2

2C

aO

*66á9

566á9

266á9

462á4

062á3

663á1

062á6

266á0

766á8

066á4

460á0

960á9

2M

nO

*0á2

50á1

20á1

90á5

11á3

60á4

90á7

90á0

00á6

10á3

10á4

60á0

0S

um

98á1

199á1

199á1

598á8

699á7

098á7

799á2

199á2

099á6

3M

gC

O3

35á1

435á9

335á5

340á1

139á4

640á1

339á9

036á8

635á7

736á3

242á9

742á6

3F

eC

O3

0á0

30á1

30á0

80á3

20á0

00á0

00á1

10á0

00á0

00á0

00á0

00á1

9C

aC

O3

64á6

063á8

464á2

259á1

459á3

659á4

559á3

163á1

463á7

063á4

256á6

357á1

9M

nC

O3

0á2

20á1

00á1

60á4

41á1

80á4

20á6

80á0

00á5

30á2

60á3

90á0

0S

um

100á0

0100á0

0100á0

0100á0

0100á0

0100á0

0100á0

0100á0

0100á0

0M

g/C

a0á4

60á4

80á4

70á5

80á5

70á5

80á5

80á5

00á4

80á4

90á6

40á6

4

*M

easu

red

valu

es;

Calc

ula

ted

valu

es;

w,

An

aly

sis

poin

tsas

mark

ed

on

Fig

s3

an

d4.



1976; Veizer et al., 1992a,b). In the Holocene,several episodes in which dolomite formed as adirect precipitate from lake water have beenreported from Lake Walyungup, Western Australia(Coshell et al., 1998), and the Coorong dolomitefrom South Australia has also been interpreted as adirect primary precipitate from lake water (Von derBorch, 1965; Rosen et al., 1988, 1989).

Most magnesite forms as a minor mineralduring diagenesis in a hypersaline environment.However, in places, hydromagnesite occurs as a

primary precipitate, i.e. Salda GoÈluÈ , south-west-ern Turkey (Braithwaite & Zedef, 1994, 1996).

Petrographic data obtained from the TF suggestthat the micritic magnesite could only haveoriginated during early diagenesis if dolomite isa primary precipitate. The micritic magnesitereplaced dolomite before the major portion ofthe burial carbonate cements formed. However,the crystalline and coarsely crystalline neomor-phic magnesite that replaced micritic magnesitegrew during late diagenesis and metamorphism.

Table 4. Carbon and oxygen isotopeanalyses of composite dolomite±magnesite samples using selectiveacid extraction.

Yield (%) d13C d18O Yield (%) d13C d18O

Depth (m/sample) Dolomite Magnesite

Magnesite Bed 24699-537á5* 11á2 21á04699-537á5 24 11á2 21á7 76 11á6 20á84699-537á6 36 11á4 22á3 64 11á5 20á64699-537á7 33 11á4 22á6 67 11á6 20á84699-537á8 24 11á3 22á7 76 11á6 21á04699-537á9 38 11á3 22á1 62 11á5 20á8

Magnesite Bed 55177-553á5* 9á0 23á15177-553á5 0 ± ± 100 9á2 23á25177-553á55 10 9á0 23á4 90 9á2 23á75177-553á6 0 ± ± 100 9á2 23á15177-553á6 0 ± ± 100 9á1 23á65177-553á65 10 7á8 22á1 90 9á0 23á65177-553á7 6 9á9 26á3 94 9á3 23á85177-553á8 0 ± ± 100 9á3 23á95177-553á9 6 8á8 22á0 74 9á1 23á7

Magnesite Bed 45177-568á5* 9á4 24á35177-568á5 9 9á2 21á8 91 9á8 24á65177-568á7 18 9á7 23á2 82 9á8 25á75177-568á8 15 9á3 22á0 85 9á8 25á45177-568á9 10á5 9á6 22á0 89á5 9á7 24á8

Magnesite Bed 35177-598á7* 9á6 21á45177-598á7 21 9á0 20á7 79 9á7 20á35177-598á7 42 9á6 23á2 58 9á6 21á85177-598á75 0 ± ± 100 9á5 21á05177-598á8 24 9á3 20á6 76 9á5 20á85177-598á85 22á5 7á1 17á4 77á5 9á5 21á85177-598á9 14 7á5 18á1 86 9á0 20á05177-598á95 21 9á3 20á8 79 9á6 21á2

Magnesite Bed 15177-799á0* 11á3 20á95177-799á0 26 11á1 20á7 74 11á5 20á15177-799á2 16 11á5 21á2 84 11á6 21á15177-799á4 18 11á3 20á8 82 11á5 20á35177-799á8 87 11á6 19á6 13 11á0 21á3

Yields are expressed as percentages of CO2 released.*Whole rock, non-selective analyses of dolomite±magnesite samples.



Several geochemical screening methods areavailable to assess the degree of diagenetic andmetamorphic alteration. Hudson (1977) foundthat oxygen isotopes may be a sensitive indicatorof diagenetic alteration. Diagenesis commonlydecreases d18O, and the effect of diagenesis canbe revealed on a d13C and d18O cross-plot. Oxygenisotopes are commonly much more easily affectedby exchangeable oxygen derived from eithermeteoric water or interstitial ¯uids at elevatedtemperatures (e.g. Fairchild et al., 1990), whereasd13C may be buffered by the pre-existing carbon-ate. In general, depletion in both oxygen andcarbon isotope values may be considerable duringlate diagenesis as well as in the course of low-grade metamorphism accompanied by deforma-tion (Guerrera et al., 1997).

Although a d13C±d18O cross-plot (Fig. 6) showsno reliable correlation, the wide spread in d18Ovalues is evident and may re¯ect resetting ofoxygen isotopes during later recrystallization. Ifthe highest d18O values of 26á3& for dolomite and25á7& for magnesite are considered as the leastaltered, then all the lower values could have beenaffected by later diagenetic and metamorphicresetting. A limited spread in d13C values suggeststhat carbon isotopes may have been buffered by

pre-existing carbonate. It is very unlikely that¯uid±rock interactions during the course of meta-morphism substantially affected carbon and oxy-gen isotope systems, as oxygen and carbon isotopevalues are not signi®cantly depleted. However,Mn enrichment in the structureless crystallinemagnesite could indicate the incorporation oflater diagenetic and metamorphic Mn-rich ¯uids.

ORIGIN OF Mg-RICH CARBONATESFROM SEDIMENTARY ENVIRONMENTS:GENERAL

Depositional environments

Typical sedimentary magnesites are con®ned to:(i) ancient marine platform carbonates (`Veitschtype'; Pohl, 1989); (ii) lacustrine sediments nearto or overlying ultrama®c rocks (e.g. Bela Stena,Nevada; Fallick et al., 1991); (iii) marine evapor-ates (e.g. Sebkha el Melah, east coast of Tunisia;Perthuisot, 1980); (iv) recent coastal salt ¯ats inarid regions (e.g. sabkhas of Abu Dhabi; Bush,1973); and (v) continental and coastal lakes (e.g.playas in Coorong Lagoon area, South Australia;Walter et al., 1973; Schroll, 1989).

Fig. 6. Plot of d13C vs. d18C comparing the TF results with magnesites from other deposits. Data are from: CoorongLagoon, South Australia (Zachmann, 1989); Lagoon, Adelaide, South Australia (Botz & von der Borch, 1984; Schrollet al., 1986); magnesite deposits of Yugoslavia (Fallick et al., 1991); Servia sedimentary magnesites (Kralik et al.,1989); Eugui, Spain, Carboniferous, coarse-grained, spar magnesite (Kralik & Hoefs, 1978); Adelaide Syncline,Copper Claim, Australia, Neoproterozoic, ®ne-grained, banded magnesite (Lambert et al., 1984); Rum Jungle, Nor-thern Territory, Australia, Palaeoproterozoic, coarse-grained, spar magnesite (Aharon, 1988); Barton, Zimbabwe,Archaean, ®ne-grained, banded magnesite (Perry & Tan, 1972; Schidlowski et al., 1975).



Most magnesite forms as a minor mineralduring diagenesis in a hypersaline environment.In places, hydromagnesite occurs as a primaryprecipitate associated with active stromatolites,i.e. Salda GoÈluÈ , south-western Turkey (Braithwa-ite & Zedef, 1994, 1996). The association ofstromatolites and both diagenetic and primarymagnesite is very common and has been reportedfrom many recent alkaline lakes (Walter et al.,1973; Last & De Deckker, 1990; Renaut, 1993;Coshell et al., 1998).

Source of Mg-rich solutions

In general, the origin of Mg-bearing solutionsforming sedimentary magnesites is not wellunderstood, although some authors suggest thatit does not form as a primary phase under surfaceconditions (e.g. MoÈller, 1989). The formation ofmagnesite requires high Mg2+/Ca2+ ratios insolution (MuÈ ller et al., 1972; MoÈller, 1989). It iswell known that Mg2+ is enriched in sea waterduring carbonate sedimentation. This processenhances the dolomitization of carbonate muds(Carpenter, 1980) and may proceed towardsmagnesitization. The in¯uence of algae in produ-cing high pH-values in the water may be animportant factor for magnesite formation; majorfossil magnesite deposits are intimately associ-ated with biohermal stromatolitic dolomite (Misra& Valdiya, 1961; Valdiya, 1969, 1995; Raha, 1980;Shevelev et al., 1991; Joshi et al., 1993). Thesalinity of the diagenetic pore solutions is veryimportant, as rising salinity enlarges the stability®eld of magnesite compared with that of dolo-mite. The Mg/Ca ratio is raised even more ifsulphates crystallize (Bathurst, 1975), leading tothe development of dense, Mg-enriched brines.

Major processes controlling oxygenand carbon isotope compositions

Carbon and oxygen isotope measurements from74 magnesite occurrences were reviewed byKralik et al. (1989). These data demonstrate thatcryptocrystalline±microcrystalline magnesites inultrama®c complexes are characterized by lowd13C values (±6& to ±18&) and relatively highd18O values (+22& to +29&). The negative d13Cand relatively high d18O values indicate theformation of magnesite at low temperatures witha meteoric carbon source (Kralik et al., 1989). Astudy of magnesite associated with ultrama®crocks of Yugoslavia by Fallick et al. (1991) sug-gested a carbon source derived from decarboxy-

lation of organic material for magnesite depositswith d13C < ±10&.

The ®ne-grained Quaternary to Recent magne-sites, which occur in evaporitic sabkha and localpond environments, exhibit relatively high d13C(+1á7& to +4á6&) and d18O values (+32& to+38&). Kralik et al. (1989) reported that thesevalues suggest magnesite precipitation in equi-librium with a dissolved inorganic carbon pool ofambient basinal water, with the generation ofisotopically very heavy carbon by fermentation.

Ancient ®ne-grained magnesites show twomain maxima d13C (+2& to +3&), d18O (+25&to +27&) and d13C (±2& to ±6&), d18O (+18& to+22&). The 13C enrichments are interpreted toresult from carbon inheritance from an evaporiticcarbonate precursor (e.g. Kralik & Hoefs, 1978). Amixture of evaporitic and meteoric waters hasbeen suggested as a parent ¯uid for the ®ne-grained magnesite with negative carbon and lowd18O (Kralik et al., 1989).

The coarse-grained spar magnesites exhibit awide range of d13C (±7á5& to +4&) and d18O (+6&to +25&). These values partly correlate with thedegree of metamorphism, without evidence ofmagnesite formation before recrystallization(Kralik et al., 1989).

Carbon and oxygen isotope compositions ofsedimentary-hosted magnesites from the maintypes of deposits are plotted in Fig. 6 in compar-ison with the TF magnesite.

ORIGIN OF Mg-RICH CARBONATES:APPLICATION TO THE TF MAGNESITE

Mechanism of 13C enrichment

The TF dolomite and magnesite exhibit extreme13C enrichment. The formation of 13C-rich dolo-stones has been ascribed to the global 2á4±2á06 Gapositive excursion of carbonate 13C/12C associ-ated with enhanced accumulation of Corg (Baker &Fallick, 1989; Yudovich et al., 1991; Karhu, 1993;Melezhik et al., 1999). A detailed study of the TFdolostones by Melezhik et al. (1999) suggestedthat, although the formation of the TF 13C-richcarbonates was driven by global factors, thecomplementary organic carbon was buried in anexternal basin. However, it has also been reportedthat TF dolostones reveal the greatest enrichmentin 13C (d13C up to +18&; Yudovich et al., 1991)known from this interval. Such enrichmentexceeds the global value for the isotopic shift at»2á0 Ga (perhaps at around +5&; Melezhik et al.,



1997, 1999). If the global d13C value of +5& wascaused by enhanced Corg burial, the furtherenrichment requires an extra 13C-rich source(s).Development of abundant stromatolite-formingmicrobial communities in shallow-water basins,establishment of evaporative and partly restrictedenvironments, high bioproductivity, enhanceduptake of 12C and penecontemporaneous recyc-ling of organic material in cyanobacterial matswith the production and consequent loss of CO2

(and CH4?) have been suggested to be additionalfactors that may have increased d13C from +5& upto +18& (Melezhik et al., 1999).

The d13C and d18O values of the TF magnesite donot differ signi®cantly from those reported for theassociated TF dolostones. Thus, we assume thatboth magnesite and dolomite obtained their carbonfrom the same reservoir and a similar mechanismwas involved in the carbon isotope fractionation.Therefore, the 13C-rich nature of the carbon sourcecan be assigned to both (i) the global carbonreservoir at 2á0 Ga and (ii) the local reservoir. Atthis stage, however, a quantitative separationbetween the local and global carbon reservoirs isnot possible. If the global d13C value is assumed tohave been around +5& (e.g. Melezhik et al., 1997,1999), then the d13C of +9á0& to +11á6& can beexplained as resulting from local factors, such asevaporation and restricted environments withhigh bioproductivity.

Depositional environments and the sourceof Mg-rich solutions

As discussed previously, the depositional set-tings of the major magnesite units (Beds 3±5)appear to have been either sabkha to playaenvironments or ponds in an upper tidal ¯at.

Epigenetic formation of magnesite seemsunlikely, as carbon and oxygen isotopes and Sr(normalized against Ca) abundances are verysimilar to those of the early diagenetic/sedimen-tary host dolomite. An external source of Mg2+-rich solutions is also unlikely, as no ultrama®crocks are documented in the area. Furthermore,lacustrine magnesites formed from ultrama®c-derived Mg solutions exhibit a large spread inboth oxygen and carbon isotope values, with thebulk of d13C being negative (Fig. 6). Therefore, seawater is the most probable source of Mg2+-richsolutions.

The Mg2+-bearing solutions were apparentlycreated in a hypersaline environment, as indica-ted by the sedimentological data. Several stages ofMg/Ca enrichment might have taken place in the

basin. First, Mg/Ca may have been enriched in seawater in the course of calcium carbonate precipi-tation in a partly closed marine environment.Secondly, an arid climate may have led to furtherevaporation of sea water-derived brine as a resultof gypsum, and then halite, precipitation. Pro-gressive evaporation and salinity increases can betraced by the development of desiccation cracksand gypsum casts in the lithofacies X dolostonesand then by halite casts in lithofacies III sedi-ments, which overly the major bed of magnesite(Bed 5, Fig. 2). With rising salinity, if calciumsulphate was precipitated, the Mg/Ca ratio wouldhave increased. The succession of these processeswould have led to the development of dense,Mg2+-enriched residual brines. Depletion of mag-nesite in Na and F compared with dolomite mayindicate de®ciency in these components in mag-nesite-forming solutions, which is consistentwith the proposition that magnesite formed fromevolved brines that lost Ca and Na by earlierprecipitation. Such Mg2+-rich brines would mostprobably have percolated downwards in thesedimentary sequence as well as laterally to thecentre of the basin, causing magnesitization.Magnesite replaced dolomite, and d13C and Srconcentration would have been buffered by thedolomite precursor, as the carbon isotope com-position as well as the Ca/Sr ratios of magnesite-bearing rocks are indistinguishable from those ofdolomite. Petrographic data suggest that themicritic magnesite could only have originatedduring early diagenesis if the dolomite was aprimary precipitate and replaced dolomite beforethe major phase of burial diagenesis. However,the crystalline and coarsely crystalline, neomor-phic magnesite, which replaced micritic magnes-ite, grew during late diagenesis and even duringmetamorphism.

Relatively increased d18O values of the magne-sites compared with the d18O average of TFdolomite are consistent with evolved evaporitic¯uids. However, the original d18O values of bothdolomite and magnesite are assumed to have beenpartially overprinted, and the magnesite isenriched in Mn. The higher d18O values andenhanced Mn concentration may be reconciled ifa sea water-derived brine was mixed with Mn-rich, 18O-depleted meteoric ¯uids. Alternatively,the enrichment in Mn and overprinting of d18Ovalues could be related to a much later stagewhen collapse breccia formed as a result ofsulphate (and halite) dissolution.

Given the limited development of magnesite inthe TF, either sabkha magnesite (similar to Abu



Dhabi) or playa magnesite (similar to the CoorongLagoon area in South Australia and Lake Wal-yungup in Western Australia) match availablesedimentological data. The d18O values of Coo-rong magnesite are similar to those of the TF,although they differ in d13C (the magnesitesstudied here are richer in 13C; Fig. 6). However,both magnesites and dolostones of the TF areenriched in 13C; part of a more general problemrelated to the global »2á0 Ga positive excursion ofcarbonate carbon.

CONCLUSIONS

1. Magnesite forms a series of 1- to 15-m-thickbeds within the »2á0 Ga Tulomozerskaya Forma-tion. The 680-m-thick unit is composed of astromatolite±dolomite±`red bed' sequence formedin a complex combination of shallow-marine andnon-marine, evaporitic, partly restricted environ-ments.

2. Dolomite-collapse breccia, stromatolitic andmicritic dolostones and sparry allochemicaldolostones are the principal rocks hosting themagnesite beds. All dolomite lithologies haveenriched d13C values of +7á1& to +11á6& andd18O ranging from 17á4& to 26á3&.

3. Magnesite occurs in different forms: struc-tureless micritic, crystalline, coarsely crystallineand ®nely laminated micritic, stromatolitic mag-nesite. All varieties exhibit positive, highlyanomalous d13C values ranging from +9á0&to +11á6& and d18O values of 20á0±25á7&.

4. Micritic magnesite originated during earlydiagenesis and replaced dolomite before themajor phase of burial. Crystalline and coarselycrystalline, neomorphic magnesite, whichreplaced micritic magnesite, formed during latediagenesis/metamorphism. Magnesite apparentlyprecipitated from sea water-derived brine, per-haps diluted by meteoric ¯uids. Magnesitizationwas accomplished under evaporitic conditions(sabkha to playa lake environment), similar to theCoorong or Lake Walyungup coastal playa mag-nesite.

5. Extremely high d13C values of magnesite anddolostones probably re¯ect a combined contribu-tion from both global and local carbon reservoirs.A 13C-rich global carbon reservoir (d13C at around+5&) is related to a perturbation of the carboncycle at 2á0 Ga, whereas the local enhancement in13C (up to +12&) was associated with evaporativeand restricted environments with high biopro-ductivity.

ACKNOWLEDGEMENTS

The results were obtained within the interna-tional project INTAS-RFBR 095-928 entitled`World-wide 2 billion-year-old isotopically heavycarbonate carbon: the evolutionary signi®canceand driving forces'. This research has beencarried out by the Geological Survey of Norway(NGU), Trondheim, jointly with the ScottishUniversities Environmental Research Centre(SUERC), Glasgow, Scotland, and the Institute ofGeology (IG) of the Russian Academy of Sciences,Petrozavodsk, Russia. Access to core material ofthe Nevskaya and Karelian Geological Expedi-tions is acknowledged with thanks. The ®eldwork was ®nancially supported by Norsk Hydro.The isotope analyses were performed at SUERCsupported by the Consortium of Scottish Univer-sities and the Natural Environment ResearchCouncil. XRF analyses were performed at NGUand ®nanced by the Kola Mineral ResourceProject. Electron-probe microanalyses and elec-tron microscope studies were carried out at theInstitute of the Continental Shelf in Trondheim.IG and, partly, NGU and SUERC were supportedby INTAS-RFBR 095-928.

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5Manuscript received 16 April 1999;revision accepted 10 August 2000.



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Palaeoproterozoic magnesite: lithological and isotopic