Stratotype of the Lower Riphean, the Burzyan Group of the Southern Urals: Lithostratigraphy,...

574

ISSN 0869-5938, Stratigraphy and Geological Correlation, 2009, Vol. 17, No. 6, pp. 574–601. © Pleiades Publishing, Ltd., 2009.Original Russian Text © M.A. Semikhatov, A.B. Kuznetsov, A.V. Maslov, I.M. Gorokhov, G.V. Ovchinnikova, 2009, published in Stratigrafiya. Geologicheskaya Korrelyatsiya,2009, Vol. 17, No. 6, pp. 17–45.

INTRODUCTION

The long-lasting discussion on principles of subdi-viding the Proterozoic into general units gave birth totwo concepts of solving the problem, i.e., to basicallydifferent chronostratigraphic and chronometricapproaches (see reviews in Cowie and Johnson, 1985;Semikhatov, 1979, 1991, 1993; Harland et al., 1990).The former approach is based on the analysis of the onlyobjective record of the Earth’s history inferable from theobservable succession of rocks and paleontologicalremains they contain and determines the boundaries ofsubdivisions considering the biotic and/or abiotic eventsin the stratigraphic record of the planet, which are recog-nizable most clearly in the type sections selected pur-posefully. Precisely this concept of applying the leadingprinciples of the general stratigraphic subdivision of thePhanerozoic to the Precambrian is the basic one in thegeneral scales of the Proterozoic subdivision, which havebeen considered during the last three decades at repre-

sentative conferences in the former Soviet Union and inthe Russia and approved by the Interdepartmental Strati-graphic Committee of the country concerned (Keller et al.,1977; Keller and Krats, 1979;

Resolution

…, 1979; Semi-khatov, 1979; Semikhatov et al., 1991;

Resolution

…,2001).

In the chronometric concept, the general subdivi-sion of the Proterozoic is based not on the geologicalrecord in its reality but on the abstract geological timeexpressed in terms of isotopic-geochronological units(Plumb and James, 1986; Harland et al., 1990; Plumb,1991; Gradstein et al., 2004a, 2004b). In this case,stratigraphic subdivisions have no links with the pale-ontological record of the Earth, are deprived of strato-types, and their boundaries correspond solely toselected isotopic dates (global standards of strati-graphic age) inferable from the analysis of the compos-ite geological record of the Earth (Plumb and James,1986; Plumb, 1991).

Stratotype of the Lower Riphean, the Burzyan Group of the Southern Urals: Lithostratigraphy, Paleontology,

Geochronology, Sr- and C-Isotopic Characteristics of Its Carbonate Rocks

M. A. Semikhatov

a

, A. B. Kuznetsov

b

, A. V. Maslov

c

, I. M. Gorokhov

b

, and G. V. Ovchinnikova

c

a

Geological Institute, Russian Academy of Sciences, Moscow, Russia;e-mail: [email protected]

b

Institute of Precambrian Geology and Geochronology, St. Petersburg, Russia

c

Institute of Geology and Geochemistry, Uralian Division, Russian Academy of Sciences, Yekaterinburg, Russia

Received December 10, 2008; in final form, January 29, 2009

Abstract

—The main objective of this work is the generalization of lithostratigraphic, biostratigraphic and iso-topic–geochronological data characterizing carbonate rocks from type succession of the broadly acknowledgedchronostratigraphic subdivision of the Lower Riphean, such as the Burzyan Group of the Southern Urals andits analogs. Using an original approach to investigation of the Rb–Sr and Pb–Pb isotopic systems in carbonatesand strict criteria of their retentivity, we studied the least altered (“best”) samples of the Burzyan carbonates,which retain the

87

Sr/

86

Sr ratio of the sedimentation environment. As long ago as 1550

±

30 and 1430

±

30 Ma,that ratio corresponded to 0.70460–0.70480 and 0.70456–0.70481. The results confirm the influx of the mantlematerial predominantly into the World Ocean of the Early Riphean. The influence of meteoric diagenesis waslikely responsible for local declines of

δ

18

O in the Burzyan carbonates down to the values of –2.5 to –1.5‰V-PDB. In the “best” samples, this parameter ranges from –0.7 to 0‰, which is consistent with the assumptionthat

δ

18

O values (0

±

1‰) characterized the stasis of the carbonate carbon isotopic composition in oceanic water2.06–1.25 Ga ago. C-isotopic data on carbonate from the Paleoproterozoic–Lower Riphean boundary forma-tions of the Urals, India, North America and Siberia suggest that the mentioned stasis ended by the commence-ment of the Early Riphean ca. 1.6–1.5 Ga ago. In the least altered carbonates of the Early Riphean, the

δ

18

Ovariation range corresponds to 4.0–4.5‰.

Key words

: Southern Urals, Lower Riphean, Burzyan Group, Sr- and C-isotopic chemostratigraphy, and the Pb–Pbage of carbonate rocks.

DOI:

10.1134/S0869593809060021

STRATIGRAPHY AND GEOLOGICAL CORRELATION

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STRATOTYPE OF THE LOWER RIPHEAN, THE BURZYAN GROUP OF THE SOUTHERN URALS 575

The main disadvantages of the chronometric scaleare as follows: (1) the subdivisions of the scale and theirboundaries are independent of the real successions ofrocks and fossils; (2) the boundaries of general subdivi-sions are defined solely in terms of the isotopic-geo-chronological units based on averaging the age valuesfor the subdivisions they separate but not on dating thereference (standard) geologic objects; (3) the chrono-metric units cannot be used for the natural periodizationof the Earth’s geological history. Nevertheless, thechronometric concept of subdividing the Precambrianbecame popular in recent years, as the InternationalSubcommission on Precambrian Stratigraphy adoptedexactly this approach in constructing the Precambriantime scale (Plumb and James, 1986) that received theinternational status recently (Plumb, 1991; Gradsteinet al., 2004a, 2004b).

An extensive multidisciplinary characterization ofthe Lower, Middle, Upper Riphean and Vendian suc-cessions, i.e., of the Precambrian subdivisions that arethe cornerstones of the Proterozoic chronostratigraphicscale, whose efficiency is well proved in Russia, is ofprime importance in the outlined situation. Thisapproach, ensuring the cross check of results obtained

by different methods when distinguishing the generalchronostratigraphic units, opens the way to use theinterchangeability principle with respect to characteris-tics of stratigraphic subdivisions in the case of their cor-relation. In this work, we consider lithostratigraphy,and the biostratigraphic and isotopic-geochronologicalcharacteristics of the Lower Riphean type successioncorresponding in Russia to the Burzyan Group of theSouthern Urals and its equivalents. In addition, the sig-nificance of original C- and Sr-isotopic data on carbon-ate rocks of the group, which are presented in the work,is evaluated in the interregional aspect.

LITHOSTRATIGRAPHY OF THE BURZYAN GROUP

The Burzyan Group distinguished in the westernflank of the Southern Urals is a basal one in the type suc-cession of the Riphean Eonothem and represents the stra-totype of the Lower Riphean Erathem. The rocks of thegroup are observable in the northeast and in the centralparts of the Bashkirian meganticlinorium, while theircomplete successions are exposed in the Taratash andYamantau anticlinoriums (Fig. 1). In the western and

Yuryuzan R.

Uph

a R.

UFA

Sim R

.

Zilim

R.

Inzer

Belaya R.

Verkhnii AvzyanMagnitogorsk

Bakal

Satka

Kusa

Zlatoust

Minyar

1

2345

6

7

8

9

100 km

55

°

N

60

°

E

1 2 3 4 5 63

Fig. 1.

Localities of the Burzyan Group sections studied in the Bashkirian meganticlinorium: (1) Archean–Proterozoic, Taratashcomplex; (2) Lower Riphean; (3) Middle Riphean; (4) Upper Riphean; (5) Vendian; (6) localities of studied sections; encircled num-bers denote (1) Kusa, (2) Satka, (3) Mt. Kazymovskaya, (4) Mt. Berezovaya, (5) Bakal, (6) Kartalinskaya Zapan, (7) Suran,(8) Ismakaevo, and (9) Askarovo sections.

Beloretsk

576


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SEMIKHATOV et al.

eastern structural-facies zones of the Taratash anticlino-rium, the sections of the group are not identical in certainaspects: deposits of the western zone show a lesserdegree of secondary alterations, are less deformed, andvolcanic rocks appear in the group’s lower part, com-posed of siliciclastic sediments, whereas in the easternzone, the abundance of carbonate members replaced bysiliciclastic deposits and the total thickness of the groupare decreased. In the western zone, the deposits experi-enced deep catagenesis, meteoric diagenesis and localmetasomatism, responsible for the deposition of second-ary magnesite and siderite. In the eastern zone, the rocksare altered under conditions of deep catagenesis,metagenesis, and local meteoric diagenesis, being partlymetamorphosed under conditions of the greenschistfacies along the Main Uralian Fault zone (Anfimov,1997; Maslov et al., 2001).

The lithostratigraphic characterization of theBurzyan Group presented below is based on the datafrom a series of joint works (Garan’, 1969;

Stratotype

…,1982, 1983; Anfimov, 1997; Anfimov et al., 1987;Kozlov et al., 1989; Maslov, 1997; Maslov et al., 2001,2004; Krupenin, 1999), whereas the geochemical data,characterizing the section intervals that have been sam-pled are original. Following Shvetsov (1958), we distin-guished pelitomorphic (grain size <0.01 mm), and thefine- (0.01–0.05 mm), small- (0.05–0.2 mm), medium-(0.2–0.5 mm), and coarse-grained (>0.5 mm) structuresof the studied grainy carbonates. The content of thesiliciclastic admixture (SA) and C

org

concentration in thecarbonates are quoted in weight percentages. The sam-pling intervals for isotopic study ranged from a few toseveral dozens meters depending on the rocks alterationextent and availability of outcrops. In the laboratory, thecollected rock samples were divided in two parts, oneused for petrographic examination, the other one forchemical and isotopic analyses. The wet chemicalmethod and the atomic absorption technique were usedto determine respectively the Ca and Mg or Mn and Feconcentrations in carbonate fractions of samples aftertheir dissolution in 1N HCl at the Laboratory of Chemi-cal Analytic Research of the GIN RAS.

Lower Riphean of the Taratash Anticlinorium

In the Taratash anticlinorium, the Burzyan Grouprests with an angular unconformity on the Archean–

Lower Proterozoic polymetamorphic Taratash Com-plex of crystalline rocks that experienced the terminalretrograde metamorphism of amphibolite facies

1800.8

±

2.6

Ma ago (Sindern et al., 2005, 2006). Con-glomerates and sandstones of the Zigalga Formation,the second one from below in the Middle Riphean typesuccession (Yurmata Group), overlie the rocks underconsideration with angular unconformity and constraintheir upper age limit. The Burzyan Group is subdividedinto three conformable lower-rank units (Fig. 2): thevolcanogenic-siliciclastic Ai (1700–2500 m), predomi-nantly carbonate Satka (1700–3500 m), and carbonate-shaly Bakal (1400–1600 m) formations. The carbonaterocks of the Satka and Bakal formations experiencedintense alteration with resultant origin of magnesite andsiderite deposits of the Satka and Bakal ore fields.

The Ai Formation

discordantly resting on the pre-Riphean crystalline complexes is divided in theTaratash anticlinorium western limb into three subfor-mations. The lower subformation (800–1300 m) begin-ning with inequigranular polymictic sandstones that areintercalated with subordinate conglomerate and brecciabeds is crowned by polymictic gravelstones, whichenclose trachybasalt flows and tuff beds of the samecomposition from 7 to 50 m thick, separated by thin(0.5–10 m) interlayers of siliciclastic sediments. Volca-nic rocks up to 250–300-m thick on the north of theTaratash anticlinorium are completely wedged out insome areas of the anticlinorium eastern limb. The mid-dle Ai Subformation (500–600 m) is composed of sand-stones, siltstones, carbonaceous shales, and rare con-glomerate and gravelstone beds, which are intercalatedwith dolomite and limestone interlayers near the unitbase. Rocks of the upper subformation (400–600 m) arecarbonaceous shales with siltstone and rare sandstoneinterlayers.

The Satka Formation

is the initial one in the car-bonate succession of the Burzyan Group and includesthe Lower and Upper Kusa, Polovinka, and Lower andUpper Satka subformations. In the general successionof the Satka Formation there are two lithostratigraphicsubdivisions that are remarkable due to their sharplydifferent composition: the Polovinka Subformation(160–200 m) in the middle part, which is composed offine-grained siliciclastic sediments containing impuri-ties of dispersed carbonaceous material, and the Kazy-movskaya Member of pure limestones (32–140 m)

Fig. 2.

Lithostratigraphy of the Burzyan Group upper part in the Taratash and Yamantau anticlinoriums and sampling levels of car-bonates for isotopic investigation: (1) bedded, (2) stromatolitic and (3) and clay limestones; (4) crystalline and (5) stromatoliticdolomites; (6) dolomite with siliciclastic admixture and (7) with phosphorite lentils; (8) low-carbonaceous shale; (9) siltstone;(10) sandstone; (11) gravelstones and conglomerates; (12) basic volcanics; (13) silicic volcanics; (14) gabbro-diabase; (15) rapakivigranite; (16) nepheline syenite; (17) stratigraphic unconformity; (18) U–Pb dates of zircons, Ma; (19) Pb–Pb age of limestones, Ma.Abbreviation of a group or formation: (Yurm) Yurmata; (Bol.In) Bolshoi Inzer; (Zg) Zigalga; (Z-K) Zigaza–Komarovo. Subforma-tion indices in the Satka and Bakal formations: (St1) Lower Kusa; (St2) Upper Kusa; (St3) Polovinka; (St4) Lower Satka;(St5) Upper Satka; (Sr1) Minyak; (Sr2) Berdagulovo; (Sr3) Angastak; (Sr4) Serdauk; (Sr5) Lapyshta. Member indices in the UpperSatka Subformation: (St5-1) Kamennaya; (St5-2) Karagai; (St5-3) Kazymovskaya. Members of the Upper Bakal Subformation:(1) Berezovaya; (2) Irkuskan; (3) Shuida; (4) Upper Shuida; (5) Gaevskii; (6) Upper Gaevskii; (7) Shikhan; (8) Upper Shikhan;(9) Upper Bakal; (10) Bulandikha.


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2009


Bu

rz

ya

n

Yur

mG

roup

Ai

Satk

aB

akal

Zg

Z-K

Form

atio

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St1

St2

St3

St5

Low

erU

pper

Subf

orm

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St5

-1S

t5-2

St5

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å‡Í

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‡fl

109876

5

4

3

2

1

St4

Form

atio

n

Subf

orm

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Zg

Z-K

Taratashanticlinorium

Yamantau anticlinorium

Bol

.In.

Sura

nY

usha

Mas

hak

1385 ± 1.4

1430 ± 30

1550 ± 30

1353 ± 5.31370 ± 51372 ± 121398 ± 28

1615 ± 45

1372±16

1

2

3

4

5

6

7

8

9

10

11

12

13

14

15

16

17

18

19

404-13

X-12 404-11404-10

404-9 404-7404-4

CI-2DCI-1D

UB-101UB-121 UB-136UB-209UB-198

UB-65UB-76

UB-81

2-21, UC-79UC-782-16, 2-18UC-72, UC-73, UC-74

UC-712-52-3

2-1

UC-55UC-54UC-53

UC-52UC-51UC-50

Sr6-45

Sr6-44

UC-45UC-44

UC-43UC-42UC-41UC-40

UC-33UC-37

UC-28UC-27

UC-26UC-24, 25UC-21, 22, 23

UC-20

UC-19UC-18

UC-17UC-16UC-15UC-9, 10, 11, 12, 13, 14UC-6, 7, 8

UC-4, 5UC3, 3A

UC-2UC-1

Sr1

Sr2

Sr3

Sr4

UB-62

Sr5

K3C-22K3C-19, 20

K3C-18K3C-16 K3C-15K3C-13, 14K3C-12K3C-9K3C-3, 4, 5, 6, 8K3C-2, K3C-2c

Mem

ber

500 m

0

UC-62UC-61

UC-6012-3

12-13

1615±45

1430±30

578


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SEMIKHATOV et al.

crowning the formation. The other subdivisions arecomposed of prevailing dolomites and subordinatelimestones. These rocks are irregularly silicified andpartially dedolomitized.

The Lower Kusa Subformation

that has been stud-ied nearby the town Kusa consists of three members:the lower and upper ones are dolomitic in composition,separated by the middle member composed predomi-nantly of limestones. Dolomites of the lower member(~170 m) are gray massive or horizontally bedded fine-grained rocks with Mg/Ca ratio ranging from 0.547 to0.657 (Table 1). The rocks enclose rare interlayers ofcross-laminated and stromatolitic dolomites, flakestonelentils, and show an irregular SA distribution (from 6.0to 68.7% near the member’s base and from 11.7 to 22.8%higher in the section). Their other characteristic featuresare a lowered Mn concentration (60–210 ppm) and rap-idly growing Fe concentration from 1490–1720 ppm inthe basal horizons to 8840 ppm at the member’s top. Atthe subformation base, there is a locally observable10-m-thick dolomite horizon with phosphorite concre-tions and lentils (Petrov et al., 1995).

The middle member (ca. 250 m) of the Lower KusaSubformation is composed predominantly of fine-grained to pelitomorphic limestones (Mg/Ca 0.004–0.055) intercalated in the lower part with rare thin inter-layers of magnesian varieties (Mg/Ca 0.092–0.121) andlimy dolomite (Mg/Ca 0.455). In the carbonate rocks,there are thin (5–15 cm) laminae of black carbonaceousand greenish gray shales and rare siltstones. The SAcontent in the rocks is high, ranging from 13.9 to 46.3%and growing in general upward in the section. Lime-stones usually contain from 110 to 280 ppm Mn,although in one sample with 42.4% SA from the lowerthird of the member, the Mn concentration is as high as540 ppm. The Fe concentration range in carbonatesfrom the middle and upper parts of the member corre-sponds to 1260–8820 ppm, being much higher (up to12300 ppm) in the calcareous sandy dolomites of thelower part, where they contain 25.3% SA.

Dolomites (Mg/Ca 0.576–0.625) prevailing in theupper member (590 m) of the Lower Kusa Subformationare similar to those of the lower member, although theyare of a lighter gray coloration. The rocks contain thin(<2–3 mm) clay and siltstone laminae, being representedby stromatolitic varieties in the lower interval 120 to130-m thick. In the dolomites, the SA content rangesfrom 7.7 to 28.8%, attaining its maximum values in themiddle of the unit, whereas the Mn concentrationincreases from 138–200 ppm near the base to 520–840 ppm close to the top. The Fe concentration in therocks changing considerably from 3170 to 12700 ppmreaches its maximum (7190–12700 ppm) in the terminalhorizons of the subformation.

The Upper Kusa Subformation

of the study areaincludes at the base a 13-m-thick packet of variegatesiltstones, shales and clay dolomites. The packet isoverlain by a succession of fine- to small-grained gray

dolomites (Mg/Ca 0.547–0.647) 130-m thick. Interca-lated with the latter rocks are interbeds of stromatoliticdolomites, chert concretions, oncolite lentils, and thin(2–5 cm) shale and occasional siltstone interlayersenriched in carbonaceous material.

The dolomites of the Upper Kusa Subformationcommonly contain from 10.0 to 20.9% SA. In places,they are enriched in fine quartz grains and grade intocalcareous sandstones (SA 42.5–51.1%) confined to theupper horizons of the subformation. The Mn concentra-tion range corresponds to 132–690 ppm in the lowerpart of the unit and to 52–175 ppm in the upper one.The distribution of Fe concentration in the rocks underconsideration is very irregular. Basal horizons of thesubformation contain 6990 to 14 300 ppm Fe, whilehigher in the section the respective values range from1270 to 4810 ppm. In addition, there are two remark-able dolomite packets, separated by depression (1130–1500 ppm) in Fe concentration: rocks of the lower onecontain 7670–10100 ppm Fe and up to 5780 ppm Fe isdetermined in the upper packet, whereas the rocks inbetween contain only 1130 to 1900 ppm Fe (Table 1).An approximately 20-m-thick member of fine-grainedlimestones (Mg/Ca 0.016–0.028) is intercalated with thedolomites in the lower third of the subformation. Rocks ofthat member contain 10.0 to 18.2% SA, 132–190 ppmMn, and 1270–1430 ppm Fe. One more 25-m-thick mem-ber of fine-grained flaggy limestones (Mg/Ca 0.009–0.033) crowns the Upper Kusa Subformation, and theserocks containing fine-grained sandy admixture (16.8–23.6%), 170–195 ppm Mn, and comparatively enrichedin Fe (2650–4030 ppm) are intercalated with shaleinterlayers, chert concretions, and flakestone lentils.

The Polovinka Subformation

(160–200 m) is com-posed of dark-colored shales depleted in carbonaceousmaterial and enclosing siltstone and rare sandstoneinterbeds. Additional marl and clay dolomite interlay-ers occur near the subformation top.

The Lower Satka Subformation

is divisible in twolower-rank units. In the vicinity of the Satka town, thelower one (200 m) is composed of gray, fine- to small-grained massive dolomites with interlayers of sandyvarieties, shales and marls. A fine-grained texture,abundance of stylolitic seams, and frequent thin vein-lets of light-colored dolomite are characteristic of dolo-mites. The rocks are enriched easterly in fine quartzgrains and grade into sandy dolomites containing up to50% SA. The upper part of the subformation(ca. 100 m) is composed of gray fine-grained clay dolo-mites and dolomitic marls with interlayers of shales andrare dolomitized limestones.

The Upper Satka Subformation

terminating sectionof the Satka Formation is divided into three members.Two lower members (Kamennaya and Karagai) arecomposed of dolomites, whereas the upper (Kazy-movskaya) member consists of limestones.

T h e K a m e n n a y a M e m b e r (130–160 m)studied near the Satka town and at the Mt. Kazy-


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Table 1.

Concentrations of trace elements, content of siliciclastic admixture (SA), C- and O-isotopic composition in carbon-ates from the Burzyan Group of the Southern Urals

Sample no.

Sampling level

1

Rock type

2

SA content

3

, %Mn, ppm

Fe, ppm

Sr, ppm Mg/Ca Mn/Sr Fe/Sr

δ

13

C V-PDB

δ

18

O V-PDB

Bakal Formation (Gaevskii Member, Upper Bakal Subformation), surroundings of the Bakal town

404-15 70 D 12.9 1460 19500 70.4 0.467 20.7 280 –2.5 –11.1

404-14 65 L 19.7 525 16300 250 0.110 2.1 64.7 –2.1 –10.5

404-13 60 L 4.6 280 3550 180 0.027 1.6 19.8 –1.2 –10.4

X-12 55 L 4.2 320 4500 230 0.011 1.4 19.6 –1.2 –10.4

404-11 50 L 6.4 190 4320 155 0.036 1.2 28.2 –1.2 –9.5

404-10 45 L 6.2 220 4730 150 0.121 1.5 31.7 –0.8 –9.4

404-9 40 L 3.7 175 2860 290 0.073 0.60 9.9 –0.7 –9.3

404-7 30 L 6.7 140 4390 110 0.079 1.3 39.9 –0.7 –9.4

404-4 20 L 7.4 280 9220 210 0.067 1.3 44.4 –1.0 –8.8

X-10 5 D 10.9 2730 38000 17 0.479 160 2240 –1.5 –10.3

Bakal Formation (Shuida Member, Upper Bakal Subformation), surroundings of the Bakal town

CI-2D 25 D 2.2 360 7080 75.6 0.556 4.6 93.7 –1.0 –14.0

CI-1D 20 D 3.3 570 9500 23.2 0.558 24.5 409.0 –0.8 –12.2

Bakal Formation (Berezovaya Member, Upper Bakal Subformation), borehole at the Mt. Berezovaya

UB-101 118 L 0.5 46 385 1090 0.003 0.04 0.4 –0.3 –8.1

UB-107 112 L 0.5 65 720 498 0.003 0.13 1.4 – –

UB-121 98 L 0.7 22 280 1080 0.003 0.02 0.3 0.0 –6.3

UB-136 84 L 0.2 15 250 1145 0.003 0.01 0.2 –0.2 –6.6

UB-153 60 L 0.7 20 230 829 0.003 0.02 0.3 – –

UB-198 58 L 0.6 16 250 1005 0.003 0.02 0.2 –0.1 –6.0

UB-209 47 L 0.2 20 260 1060 0.003 0.02 0.2 0.0 –6.1

UB-62 30 L 0.5 19 210 765 0.003 0.02 0.3 –0.3 –7.0

UB-65 25 L 0.7 22 210 830 0.003 0.03 0.3 –0.2 –7.4

UB-73 17 L 0.5 20 230 1102 0.003 0.02 0.2 – –

UB-76 13 L 0.5 25 250 975 0.003 0.03 0.3 –0.3 –7.6

UB-81 6 L 0.7 40 390 810 0.003 0.05 0.5 –0.3 –7.6

Satka Formation (Kazymovskaya Member, Upper Satka Subformation), Mt. Kazymovskaya

2-24 30 L 3.4 40 1040 – 0.034 – – – –

2-23 29 L 2.2 62 920 2150 0.014 0.03 0.4 – –

2-22 28 L 1.4 29 810 – 0.003 – – – –

UC-79 22 L 3.5 31 790 2450 0.003 0.01 0.3 –0.4 –8.6

2-21 21 L 0.8 14 280 2740 0.005 0.01 0.1 –0.5 –7.6

UC-78 20 L 0.9 11 250 2750 0.003 0.00 0.1 0.0 –6.3

2-18 19 L 1.7 20 460 2180 0.005 0.01 0.2 –0.6 –7.9

2-16 15 L 2.2 18 700 2340 0.006 0.01 0.3 –0.4 –8.0

UC-77 13 L 4.1 30 920 1220 0.003 0.02 0.8 – –

2-14 12 L 3.7 19 630 1250 0.056 0.02 0.5 – –

UC-76 11 L 3.7 25 660 1420 0.003 0.02 0.5 – –

UC-75 10 L 4.8 35 1030 1005 0.003 0.03 1.0 – –

2-10 9 L 2.9 23 800 1140 0.002 0.02 0.7 – –

UC-74 8 L 1.1 21 420 1680 0.005 0.01 0.3 –0.1 –6.6

UC-73 7 L 3.1 22 730 1490 0.003 0.01 0.5 –0.4 –7.3

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SEMIKHATOV et al.

Table 1.

Contd.)

Sample no.

Sampling level

1

Rock type

2

SA content

3

, %Mn, ppm

Fe, ppm


δ

13

C V-PDB

δ

18

O V-PDB

UC-72 5 L 3.3 27 1120 1180 0.008 0.02 0.9 –0.5 –6.9

2-7 2 L 2.5 160 2550 – 0.003 – – – –

Satka Formation (Karagai Member, Upper Satka Subformation), Mt. Kazymovskaya

UC-71 640 D 1.3 45 1200 24.6 0.589 1.8 48.8 –0.2 –6.9

2-5 638 D 2.9 120 1740 – 0.597 – – – –

2-3 635 D 0.7 85 740 66.4 0.599 1.3 11.1 –0.1 –8.1

2-1 630 D 2.4 63 560 24.8 0.594 2.5 22.6 +0.1 –5.3

Satka Formation (Karagai Member, Upper Satka Subformation), surroundings of the Satka town

UC-55 350 D 3.8 135 710 34.2 0.637 4.0 20.8 –0.2 –6.9

UC-54 340 D 4.7 125 2500 32.8 0.625 3.8 76.2 –0.9 –3.4

UC-53 330 D 15.5 160 5750 36.4 0.626 4.4 158.0 –0.4 –3.7

UC-52 320 D 5.3 142 3450 38.1 0.614 3.7 90.6 –0.4 –3.9

UC-51 210 D 2.3 57 550 37.3 0.614 1.5 14.7 –0.4 –5.9

UC-50 205 D 1.0 73 630 39.3 0.614 1.9 16.0 –0.4 –8.1

ST6-45 10 D 1.0 70 600 44.2 0.610 1.6 13.6 –0.7 –7.1

Satka Formation (Kamennaya Member, Upper Satka Subformation), surroundings of the Satka town

ST6-44 –10 D 9.6 150 800 51.0 0.610 2.9 15.7 –0.6 –5.6

Satka Formation (Upper Kusa Subformation), surroundings of the Kusa town

UC-45 1280 L 16.8 170 2650 59.3 0.009 2.9 44.7 – –

UC-44 1275 L 23.6 195 4030 – 0.033 – – – –

UC-43 1130 D 18.3 175 5780 36.2 0.641 4.8 160 – –

UC-42 1120 D 19.4 130 4810 – 0.631 – – – –

UC-41 1115 D 51.1 57 1190 48.3 0.610 1.2 24.6 – –

UC-40 1075 D 42.5 52 1130 50.7 0.627 1.0 22.3 –0.4 –6.9

UC-39 1055 D 14.3 105 1910 – 0.617 – – – –

UC-38 1035 D 20.9 85 1920 – 0.647 – – – –

UC-37 1030 D 10.4 76 1410 31.2 0.634 2.4 45.2 –0.6 –7.0

UC-36 988 D 14.5 690 10100 – 0.611 – – – –

UC-35 985 D 18.8 250 7670 43.2 0.571 5.8 176 – –

UC-34 1080 D 27.3 305 9990 – 0.547 – – – –

UC-33 1070 L 10.0 132 1430 69.9 0.028 1.9 20.5 –0.9 –11.3

UC-32 1065 L 18.2 190 1270 – 0.016 – – – –

UC-31 1050 D 29.1 530 14300 26.1 0.586 20.4 550 – –

UC-30 1015 D 13.1 400 6990 – 0.600 – – – –

UC-29 1005 D 19.0 495 12500 30.9 0.562 16.0 403 – –

Satka Formation (Lower Kusa Subformation), surroundings of the Kusa town

UC-28 980 D 21.6 520 7190 44.9 0.589 11.6 160 – –

UC-27s 965 D 11.2 840 12700 35.8 0.580 23.3 353 – –

UC-26 960 D 15.0 280 7220 – 0.593 – – – –

UC-25 895 D 16.6 175 4410 41.0 0.621 4.3 108 – –

UC-24 880 D 12.4 180 4100 – 0.614 – – – –

UC-23 860 D 27.6 227 5790 32.2 0.590 7.1 180 – –

UC-22 850 D 28.8 230 5780 – 0.576 – – – –

UC-21 835 D 8.1 235 3840 – 0.625 – – – –


Vol. 17

No. 6

2009


Table 1.

Contd.)

Sample no.

Sampling level

1

Rock type

2

SA content

3

, %

Mn, ppm

Fe, ppm


δ

13

C V-PDB

δ

18

O V-PDB

UC-20s 830 D 13.0 165 3890 21.9 0.615 7.5 178 –1.1 –8.7

UC-19 440 D 7.7 138 3170 – 0.623 – – – –

UC-18 420 D 17.1 200 5120 35.1 0.606 5.7 146 – –

UC-17 335 L 18.3 190 4220 – 0.045 – – – –

UC-16 305 L 46.3 220 3470 – 0.011 – – – –

UC-15 290 L 13.9 205 1790 56.8 0.004 3.6 31.5 –0.4 –10.7

UC-14 260 L 25.5 225 5100 – 0.055 – – – –

UC-13 250 L 18.7 224 4950 44.9 0.092 5.0 110 – –

UC-12 238 L 39.0 180 4020 – 0.039 – – – –

UC-11 230 L 42.4 540 2540 52.3 0.051 10.4 48.8 – –

UC-10 225 L 18.1 225 8320 – 0.093 – – – –

UC-9 220 L 21.7 280 8820 55.2 0.121 5.1 160 – –

UC-8 190 D 25.3 250 12300 – 0.455 – – –

UC-7 185 L 21.1 195 8020 53.4 0.268 3.7 150 – –

UC-6 175 L 16.3 110 1260 85.4 0.016 1.3 14.8 –0.6 –10.6

UC-5 160 D 22.5 210 8840 – 0.657 – – – –

UC-4 150 D 22.8 170 6670 27.1 0.619 6.3 247 – –

UC-3A 120 D 11.7 110 2540 – 0.653 – – – –

UC-3 110 D 6.0 105 2740 39.3 0.659 2.7 69.7 – –

UC-2 45 D 68.7 60 1490 37.4 0.646 1.6 39.8 –0.6 –9.6

UC-1 15 D 23.1 105 1720 – 0.675 – – – –

Suran Formation (Lapyshta Subformation), surroundings of the Kartalinskaya Zapan village

K3C-22 450 L 11.4 80 1530 343 0.004 0.24 4.5 0.1 –13.4

K3C-21 440 L 17.1 107 2310 – 0.023 – – – –

K3C-20 420 L 15.5 160 2980 322 0.023 0.50 9.3 – –

K3C-19 400 L 26.8 165 1510 – 0.004 – – – –

K3C-18 345 L 4.6 114 1250 368 0.008 0.31 3.4 –0.5 –13.4

K3C-17 335 L 8.9 108 1190 – 0.003 – – – –

K3C-16 325 L 6.0 120 1040 359 0.003 0.33 2.9 –0.6 –14.6

K3C-15 320 L 18.0 129 1290 – 0.009 – – – –

K3C-14 245 L 26.0 215 2250 321 0.020 0.67 7.0 – –

K3C-13 225 L 31.3 145 1570 – 0.010 – – – –

K3C-12 185 L 3.5 125 1230 344 0.003 0.36 3.6 –0.7 –13.8

K3C-11 160 L 20.9 117 1790 – 0.010 – – – –

K3C-10 150 L 10.2 93 1310 – 0.011 – – – –

K3C-9 134 L 7.8 80 1060 350 0.017 0.23 3.0 0.0 –11.6

K3C-8 125 L 21.0 62 910 – 0.010 – – – –

K3C-6 90 L 29.6 115 2170 344 0.033 0.33 6.3 – –

K3C-5 55 L 15.4 137 1480 – 0.003 – – – –

K3C-4 48 L 6.8 115 1470 335 0.006 0.34 4.4 – –

K3C-3 37 L 15.4 125 1820 – 0.005 – – – –

K3C-2c 15 L 0.6 120 1090 285 0.006 0.42 3.8 –0.6 –14.0

K3C-2 14.9 L 8.6 130 1170 320 0.011 0.41 3.7 –0.6 –14.1

K3C-1 10 L 10.9 144 1770 – 0.047 – – – –

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SEMIKHATOV et al.

movskaya site is composed of massive to thin-beddeddolomites with marl, shale and calcareous shale inter-layers. As in the Lower Satka Subformation, the thinnetwork of veinlets filled in by light secondary dolo-mite and stylolitic seams are observable in the dolo-mites. Judging from the only analyzed sample, thedolomites (Mg/Ca 0.610) from the Satka section con-tain 9.6% SA, 150 ppm Mn, and 800 ppm Fe near themember’s top.

T h e K a r a g a i M e m b e r. The lower thickerpart (300–400 m) of the member has been studied nearthe town of Satka, and the upper one at the Mt. Kazy-movskaya site (Fig. 1). In the lower part of the member,the rocks are represented by dark gray, small- tomedium- and coarse-grained dolomites with interbedsof their brecciated varieties and thin (up to 5–6 cm)laminae of marl and black shales. In addition, charac-teristic of dolomites are nests and numerous crosscut-ting veinlets (up to 1.5–2 cm thick) of white secondarycalcite. Bodies of carbonate breccia irregular in shapeand up to 25-m across occur in the ore field and containwedge-like segregations of white crystalline dolomiteresembling in morphology the pseudomorphs aftergypsum crystals. Accordingly, it has been suggestedthat brecciated structures were associated in origin withdissolution of sulfate minerals that existed in sediments(Krupenin and Prochaska, 2005). Crystalline magnesitedeposits of the Satka district are confined to lower partof the Karagai Member and partially to the underlyingKamennaya Member. Stratiform magnesite bodiesfrom 10 to 80-m thick are of metasomatic origin (Anfi-mov et al., 1983). Judging from composition of fluid

inclusions in magnesite and host dolomites, carbonatesof the Karagai Member from the Satka area experi-enced metasomatism under the influence of high-Mgbrines with low Cl/Br and Na/Br ratios (Krupenin andProchaska, 2005). In addition, the rocks of the KaragaiMember surround the Berdyaush massif of rapakivigranites and associated rocks.

Near the Satka town, the Mg/Ca ratio in dolomitesof the Karagai member increases upward in the sectionfrom 0.610 to 0.637, whereas the SA content is usuallynot higher than 5.3%. The SA content is as high here as15.5% only in a small rock packet confined to the sec-tion upper part, where the Mn (160 ppm) and Fe(5750 ppm) are at the maximum in the member. TheMn concentration ranges in rocks of the member from70 to 160 ppm growing in general upward in the sec-tion, whereas the Fe concentrations are within the rangeof 600–5750 ppm, and their highest values (2500–5750 ppm) are recorded in the middle and upper part ofthe member. At the Mt. Kazymovskaya site, the upperdolomites of the Karagai Member are somewhat morecalcareous (Mg/Ca 0.589–0.599) and depleted in thecomponents under consideration as compared to rocksexposed in the Satka area. The SA content ranges herefrom 0.7 to 2.9%, the Mn concentration from 45 to120 ppm, whereas the Fe concentration grows upwardin the section from 560 to 1200–1740 ppm.

T h e K a z y m o v s k a y a M e m b e r of lime-stones that is only 32-m thick at the eponymous siteincludes a basal bed (1 m) of synsedimentary carbonatebreccia and dolomitic shale. The overlying greater partof the member is composed of gray pelitomorphic and

Table 1. Contd.)

Sample no.

Sampling level1

Rock type2

SA content3, %

Mn, ppm

Fe, ppm

Sr, ppm Mg/Ca Mn/Sr Fe/Sr δ13C

V-PDBδ18O

V-PDB

Suran Formation (Lapyshta Subformation), surroundings of the Askarovo village

AC-6 220 D 14.0 2380 14600 – 0.606 – – – –

AC-5 210 D 23.2 3060 26800 – 0.590 – – – –

AC-4 130 D 23.0 1170 15900 – 0.602 – – – –

AC-3 120 D 11.0 3090 27900 – 0.595 – – – –

AC-2 110 D 12.1 3560 31200 – 0.594 – – – –

AC-1 30 D 16.4 920 17500 – 0.588 – – – –

Suran Formation (Minyak Subformation), surroundings of the Ismakaevo village

I2-3 25 L 16.6 350 6010 182 0.075 1.92 33.0 –0.6 –15.3

I2-13 20 L 4.9 160 1490 700 0.029 0.23 2.1 –0.2 –13.3

Suran Formation (Minyak Subformation), surroundings of the Suran village

UC-62 37 D 15.3 240 5950 – 0.634 – – – –

UC-61 35 D 13.0 280 6280 – 0.625 – – – –

UC-60 30 D 23.1 310 650 24.5 0.645 12.7 26.5 0.2 –13.9

Notes: (1) Sampling levels are designated in meters above the base of carbonate rocks exposed in the respective section or below the topin the case of the Kamennaya Member; (2) letters in the rock type column denote limestone (L) or dolomite (D); (3) siliciclasticadmixture; hyphen designates data absence.

STRATIGRAPHY AND GEOLOGICAL CORRELATION Vol. 17 No. 6 2009


fine-grained limestones with Mg/Ca ratio not greaterthan 0.002–0.008 in most cases except for three sam-ples, where this ratio is increased up to 0.014-0.056.The SA content in the rocks is usually within the rangeof 0.8–3.7%, being sometimes as high as 4.8–9.8%.The Mn concentration in the basal rocks corresponds to160 ppm and declines upward in the section down to11–62 ppm. The Fe concentration declines simulta-neously from 2550 to 420–1140 ppm. The Mn and Feconcentrations in the limestones of the KazymovskayaMember are the lowermost ones throughout the SatkaFormation in contrast to the highest Sr concentrations(1005–2750 ppm) in these rocks.

The Bakal Formation terminating the BurzyanGroup succession in the Taratash anticlinorium isdivided in sections of the Satka and Bakal areas intolower and upper subformations. The upper subforma-tion is overlain with angular unconformity by sand-stones of the Zigalga Formation second from below inthe Middle Riphean type succession.

The Lower Bakal Subformation (400–450 m) iscomposed of gray to black shales and phyllite-likerocks, which are intercalated with subordinate interlay-ers of calcareous shales and quartz or feldspar-quartzsiltstones.

The Upper Bakal Subformation (700–900 m)includes five members of low thickness, which arecomposed predominantly of gray carbonate rocks, thestromatolitic varieties inclusive. These are the Bere-zovaya (80–150 m), Shuida (100–120 m), Gaevskii(40–60 m), Shikhan (90–100 m) and Upper Bakal(100–200 m) members. Each of these members is over-lain in turn by members of comparable thickness,which are composed of fine-grained siliciclastic rocks(low-carbonaceous and phyllite-like shales with silt-stone interlayers). The overlying units are termed theIrkuskan (120–300 m), Upper Shuida (30–40 m),Upper Gaevskii (70 m), Upper Shikhan (60 m) andBulandikha (up to 20 m) members. In the vicinity of thetown of Bakal, carbonates of the subformation arelocally replaced by siderite that originated in the courseof the catagenetic alteration of the Riphean clayey sed-iments and penetration of Fe-bearing fluids into the car-bonate reservoirs during the tectonic activation of theregion (Krupenin, 1999) approximately 1010 ± 100 Maago (Kuznetsov et al., 2005).

Outside the ore field, the carbonate members of theUpper Bakal Subformation are composed predomi-nantly of limestones except for the Shuida and Gaevskiimembers, the former of dolomitic composition and thelatter one including two beds of dolomite. Massiverocks associated with less common stromatolitic variet-ies, flakestone interlayers and lentils usually prevail inthe lower parts of the limestone members, whereas theirmiddle parts are composed of massive limestone only,and oncolitic varieties with intraclastite lentils are con-fined to the upper parts of the members. In the ShuidaMember, basal stromatolitic dolomites with cross-lam-

inated interbeds grade upward into massive small-grained dolomites with marl and shale interbeds,whereas its terminal part is composed of fine- to small-grained dolomites enclosing interlayers of low-carbon-aceous shales, siltstones, and rare sandstones. TheBulandikha member at the formation top is representedat its base by shales and sandstones, being crowned bystromatolitic limestones. Subhorizontal bedding andprevalence of fine- to small-grained textures are charac-teristic features of the Bakal carbonates outside the orefield, whereas rosette and banded structures originatedduring the epigenetic recrystallization (Krupenin,1999) are typical of the rocks in the ore field. Westwardand especially to the east of the considered sections, thecarbonate members of the Upper Bakal Subformationbecome gradually replaced by alternating beds of feld-spar-quarts sandstones, and dark gray carbonaceousand phyllite-like shales (Stratotype …, 1983). We sam-pled carbonates for isotopic analysis only from thelower Berezovaya, Shuida and Gaevskii members,because the rocks of the Upper Shikhan and UpperBakal members experienced epigenetic alterations notonly under the influence of elision fluids derived fromsiliciclastic members of the section, but also during themeteoric diagenesis of the preZigalga time (Anfimov,1997; Krupenin, 1999; Kuznetsov et al., 2005).

T h e B e r e z o v a y a M e m b e r (80 m) is com-posed in the Mt. Berezovaya section of pelitomorphiclimestones that reveal the absolute minimum of Mgcontent among the Burzyan carbonates (Mg/Ca<0.003), a low SA percentage (0.2–0.7%), and loweredMn concentration (15–40 ppm as a rule and up to 46–65 ppm near the member’s top). The Fe concentrationcommonly ranging from 210 to 280 ppm is increasedup to 390 and 385–720 ppm near the base and top of themember, respectively. The quoted geochemical charac-teristics suggest that limestones of the member are verysuitable rocks for isotopic studies.

T h e S h u i d a M e m b e r is represented near thetown of Bakal by coarse- and small-grained calcareousdolomites (Mg/Ca 0.556–0.558), which are relativelyenriched in Mn (360–570 ppm) and Fe (7080–9500 ppm), despite the low SA content (2.2–3.3%).The rocks are intercalated with magnesite and shaleinterbeds, quartzite lentils, and stratiform to lenticularsiderite bodies.

T h e G a e v s k i i M e m b e r has been studied inthe same section along with the Shuida Member. Itincludes basal and crowning 5- to 7-m-thick beds ofmedium-grained calcareous dolomites (Mg/Ca 0.467–0.479) containing 10.9–12.0% SA, being composed inthe thicker central part (60 to 70 m) of fine-grained lime-stones that have a variable Mg/Ca ratio (0.011–0.121).These Conophyton varieties of limestones contain 3.7 to7.4% SA that is most concentrated near the member top(up to 19.7%). The Mn concentration ranges in lime-stones from 140 to 525 ppm, being lower in the middlepart of the member. The distribution of the Fe concentra-

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SEMIKHATOV et al.

tion is extremely irregular, as it corresponds to 9220 and16300 ppm at the member’s base and top, respectively,and varies from 2860 to 4730 ppm in the intermediateinterval. Similarly incomparable are the Mn and Fe con-centrations in the basal and terminal dolomite beds ofthe Gaevskii Member. The basal beds contain2730 ppm Mn and 38000 ppm Fe, whereas in the upperbed the respective values are decreased down to 1460and 19500 ppm.

Thus, pelitomorphic limestones of the Kazy-movskaya and Berezovaya members can be regarded asthe least altered rocks among the carbonates of theSatka and Bakal formation according to the lowermostMg/Ca ratio (<0.008), SA content, and Mn and Fe con-centrations established in these rocks (Table 1). Theother limestones with fine- and/or small-grained tex-tures reveal indications of epigenetic recrystallization,while the small- and coarse-grained dolomitic varietiesof carbonate rocks contain the late dolomite veinlets.

Lower Riphean of the Yamantau Anticlinorium

In the Yamantau anticlinorium, equivalents of theAi, Satka and Bakal formations are the Bolshoi Inzer,Suran and Yusha formations also separated from eachother by conformable boundaries (Stratotype …, 1983;Maslov et al., 2001). The base of the Bolshoi Inzer For-mation is unexposed, while the Yusha Formation isoverlain with angular unconformity by the Mashak For-mation of volcanogenic and siliciclastic rocks, the basalone in the Middle Riphean type succession. TheMashak Formation conformably underlies in turn theZigalga sandstones.

The Bolshoi Inzer Formation (2200 m) is com-posed of small- and medium-grained sandstones withlentils and interlayers of low-carbonaceous shales, raresmall-pebbled conglomerates, fine-grained limestonesor inequigranular dolomites enriched in clayey andsandy material.

The Suran Formation (1000–2800 m) is dividedinto the Minyak, Berdagulovo, Angastak, Serdauk, andLapyshta subformations. The lower and upper subfor-mations are of carbonate composition, whereas thethree intermediate subdivisions are composed of fine-grained siliciclastic rocks.

The Minyak Subformation is 280 to 400-m thicknear the Ismakaevo village. The basal cross-beddedlimestones (Mg/Ca 0.029–0.075) of the subformationare up to 35-m thick and contain 4.9–16.6% SA, 160–350 ppm Mn, and 1490–6010 ppm Fe. In this section,the dolomites in the upper part of the unit are replacedby crystalline magnesite. In the vicinity of the Suranvillage, the formation is 300–350-m thick, composedpredominantly of gray small-grained dolomites interca-lated with flakestone lentils and rare interlayers ofblack shales. In the lower 40-m-thick interval, the dolo-mites (Mg/Ca 0.625–0.645) are enriched in SA (13.0–

23.1%) and Mn (240–310 ppm), whereas the Fe con-centration ranges here from 650 to 6280 ppm.

The Berdagulovo (250–500 m), Angastak (250–1000 m) and Serdauk (300–1000 m) subformations arecomposed predominantly of low-carbonaceous shaleswith marl and sandstone interlayers. In addition, thereare pyrite nodules and subordinate thin siltstone andsandy dolomite interlayers in the Berdagulovo Subfor-mation, rare interbed of sandy dolomites and fine-grained sandstones in the Angastak Subformation,whereas interbeds and lentils of clay limestones in theSerdauk Subformation are enriched in dispersed car-bonaceous material. Slumping and convolute structuresare typical of rocks in the Berdagulovo Subformation.

The Lapyshta Subformation, the crowning succes-sion of the Suran Formation, has been studied in twosections of different compositions: one section of thedolomites (600–650 m) is exposed near the Askarovovillage in the south of the anticlinorium, and the otherone of limestones (~500 m) is located in the central areaof the structure close to the Kartalinskaya Zapan village(Fig. 1). In the Askarovo section, the subformation iscomposed in its lower part of gray massive to beddedsmall- and coarse-grained dolomites irregularly silici-fied and containing veinlets of white crystalline dolo-mite. In the upper part, the dolomites are intercalatedwith interlayers and lentils of oncolitic varieties. Dolo-mites (Mg/Ca 0.588–0.606) with a considerable SApercentage (11.0–23.2%) are irregularly enriched inMn and Fe. At the subformation base, the Mn concen-tration is 920 ppm, whereas higher in the section itranges from 1170 to 3560 ppm. The basal beds of theunit contain up to 17500 ppm Fe; in the middle beds,the concentration of this element varies from 15900 to31200 ppm, declining down to 14600 ppm near the top.In the middle part of the unit, there is a member (100–150 m) of dark-colored low-carbonaceous shales andsubordinate siltstones. In the Kartalinskaya Zapan sec-tion, the subformation is composed of gray small- tomedium-grained limestones, with a characteristicMg/Ca ratio ranging from 0.003 to 0.011 (occasionallyincreasing to 0.033 and 0.047). The SA content in thelimestones declines here upsection from 29.6 to 0.6%,the Mn concentration is low (62–215 ppm), and the Feconcentration grows from 910 to 2980 ppm toward thetop of the subformation. The limestones contain ratesiltstone interlayers and veinlets of epigenetic calcite.

The Yusha Formation (600–1000 m), the crowningsection of the Burzyan Group in the Yamantau anticlino-rium, is composed in its lower part (150–300 m) of low-carbonaceous shales with siltstone, sandstone, and rarelimestone interlayers. In the middle part (350–450 m),the rocks are represented by shales alternating with silt-stones and/or calcareous sandstones, whereas shales arethe dominant rocks in the upper part (200–300 m).

Deposits of the Burzyan Group accumulated underthe influence of directed changes in facies environ-ments. At the initial stage of the Ai and Bolshoi Inzer



time, disintegration products of the Early Proterozoicweathering crust (Maslov et al., 2004) accumulated ina wide spectrum of basinal settings, varying from deep-water gravitational to intertidal. During the middlestage of the Satka and Suran time, carbonate and silici-clastic material that played the leading part in sedimen-tation accumulated between the fair-weather and stormwave bases and even below the last level. This was aperiod, when accumulated shoal settings of cross-bed-ded sandy carbonates, oncolites, and stratiform stroma-tolites developed in places only. At the terminal stage ofthe Bakal and Yusha period, the basin deepened andgave rise to the expansion of stromatolites below thestorm wave base and to an accelerated influx of fine-grained siliciclastic material into the open space in theeastern part of the basin.

CHRONOMETRIC CONSTRAINTS OF THE BURZYAN GROUP

For a long time, the maximum age limit was esti-mated for the Burzyan Group as corresponding to 1650–1900 Ma, based on the first K–Ar, Rb–Sr and U–Pb iso-topic dates (the results of thermoisochron method inclu-sive) for the pre-Burzyan crystalline rocks (Garan’,1969; Garris, 1964; Salop, 1973; Stratotype …, 1983,and references in these works). As currently established,the Rb–Sr isochron age of the retrograde metamorphismunder conditions of amphibolite facies that accom-plished the transformation of the pre-Burzyan polymeta-morphic basement corresponds to 1800.8 ± 2.6 Ma(Sindern et al., 2005, 2006). The event that separated theconclusive stages of the aforementioned metamorphictransformations and first stages of sedimentation on thepassive margin of the East European craton (the BurzyanGroup and its equivalents) was dated by isotopic meth-ods at 1650 ± 50 Ma in the last decade of the 20th century(Semikhatov et al., 1991). The estimated value wasbased on the interpolation of the U–Pb isotopic dateobtained for the volcanogenic zircons from the Lower AiSubformation (1615 ± 45 Ma, Krasnobaev, 1986) and forthe youngest complexes of the Trans-Scandinavian gran-ite porphyry belt (1780–1680 Ma), interpreted as the lat-est stage of the Early Proterozoic magmatism in the des-ignated craton (Gaal and Gorbatschev, 1987).

The minimal age limit of the Burzyan deposits wasdetermined in the Taratash anticlinorium, based on theisotopic dates obtained for two intrusive bodies: the Ber-dyaush massif of rapakivi granites is localized in theKaragai Member of the Satka Formation (Kozlov et al.,1989) and the so-called Main Dyke that intruded on therocks of the Upper Bakal Subformation (Stratotype …,1983). Initial determinations of emplacement time forthese bodies resulted in controversial interpretations oftheir ages (Garan’, 1969; Garris, 1964; Salop andMurina, 1970; Salop, 1973). As it was shown later on(Krasnobaev et al., 1981, 1984), the Rb–Sr age of differ-ent rocks from the Berdyaush massif corresponds to1348 ± 13 Ma, and data points characterizing accessory

zircons from granites, quartz syenite-diorites, and thegabbro of the massif define the discordia age of 1354 ±10 Ma (Krasnobaev, 1986). The U–Pb age of 1353 ±5 Ma was established in addition for zircons from thecentral part of the feldspar ovoid sampled in the Ber-dyaush granites (Belyaev et al., 1996).

The U–Pb (SHRIMP-II) dates published recentlyfor zircons from the Berdyaush massif rocks of differ-ent petrographic composition suggest a somewhat olderage of the massif as compared to the earlier inferences(Ronkin et al., 2005, 2007). Zircons from the gabbrowere dated at 1389 ± 28 Ma, from the quartz syenite-porphyry at 1372 ± 12 Ma, from the rapakivi granites at1370 ± 5 Ma, and from the nepheline syenites at1368.4 ± 6.2 and 1373 ± 21 Ma. These values are con-sistent with the Sm–Nd age of 1371 ± 26 Ma deter-mined for whole-rock samples of the rapakivi granitesand biotite, apatite, and zircon separated from theserocks (Ronkin and Lepikhina, 2006). In addition, thebaddeleyite from the Main Dyke that is localized in theBakal Formation is dated, using the U–Pb method, at1385.3 ± 1.4 Ma (Ernst et al., 2006), while the Rb–Srage of biotite from the dyke corresponds to 1360 ±35 Ma (El’mis et al., 2000).

In the Yamantau anticlinorium, the minimal agelimit of the Burzyan Group can be inferred from isoto-pic dates obtained for volcanics sampled close to theMashak Formation base that rests with angular uncon-formity on the Lower Riphean strata and is the lower-most subdivision in the Middle Riphean type succes-sion. The least altered (devoid of sericitization) liparite-dacites of the formation contain zircons having the U–Pbage of 1370 ± 16 Ma (SHRIMP–II, Ronkin et al., 2007).Consequently, the radiometric ages characterizing ter-minal horizons of the Burzyan Group in the Taratashand Yamantau anticlinoriums appear to be concordantwithin the analytical uncertainty, and the lithostrati-graphic markers of two important regional events con-firm the correlation of these horizons. One of the eventswas transition from the accumulation of coarse silici-clastic material (deposits of the Ai and Bolshoi Inzerformations) to the deposition of thick carbonate succes-sions corresponding to the Lower Kusa Subformationof the Satka Formation and the Minyak Subformationof the Suran Formation. The second event was thereplacement of carbonate sedimentation that gave ori-gin to deposits of the Satka and Suran formation by thedeposition of shales of the Lower Bakal Subformationand fine-grained siliciclastic sediments confined tolower part of the Yusha Formation.

As already mentioned above, recent isotopic-geo-chronological dates suggest that the accumulation ofthe Burzyan Group commenced after 1800.8 ± 2.4 Ma(after termination of metamorphic transformations inthe preBurzyan basement) but before 1615 ± 45 Ma (theeruption time of the mid-Ai volcanic rocks). The eventin question cannot be dated more precisely and proba-bly was close to the eruption time of the Ai volcanics,

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as the accumulation of siliciclastic deposits underlyingthe volcanics barely took a long time. At first glance,the above assumption contradicts the K–Ar dates(1594–1329 Ma), characterizing the whole-rock sam-ples of basalts from equivalents of the Ai Formation inthe Kama–Belaya aulacogen (Stratigraphic …, 2000).However, the quoted dates are mutually discordant andcan be omitted from consideration because of the K–Arsystem lability in the whole-rock basaltic samples.

The Pb–Pb isochron dates obtained recently for lime-stones from the Kazymovskaya Member of the SatkaFormation (1550 ± 30 Ma) and from the BerezovayaMember of the Bakal Formation (1430 ± 30 Ma) firstelucidate the time of the early diagenesis in the carbonatesediments of the Burzyan Group (Kuznetsov et al., 2005,2003a, 2008). The dates are reliable in the methodicaspect, as evident from the following. (1) The quotedauthors, who studied the U–Pb systematics of samplesfrom the Kazymovskaya and Berezovaya members, reliedupon strict geochemical criteria that guarantee the reten-tivity of isotopic systems in limestones: Mn/Sr < 0.2,Fe/Sr < 5, Mg/Ca < 0.024, δ18O > –10‰ V–PDB. (2) Anisotopic-geochronological study was carried out only forthose samples, which characterize the Sr-isotopic com-position in seawater of the Early Riphean time. (3) Theused stepwise dissolution of samples in HBr secured theseparation of the carbonate fractions for analysis, whichare the least altered in terms of the U–Pb systematics andsuitable for obtaining the correct geochronological infor-mation (Ovchinnikova et al., 1998, 2000; Kuznetsovet al., 2003a, 2005, 2008). (4) The analytical resultsshowed no correlation between data points in the206Pb/204Pb–1/204Pb and 208Pb/204Pb–206Pb/204Pb dia-grams in contrast to the mixing systems and systems ofepigenetic recrystallization. In addition, the Pb–Pb datesobtained for the limestones under consideration are con-sistent, if the position of the rocks in the Burzyan succes-sion is taken into consideration, with the U–Pb, Sm–Ndand Rb–Sr dates known for the magmatic rocks of theSouthern Urals (Fig. 2).

The date of 1400 ± 10 Ma (±2σ), obtained by theU–Pb (SHRIMP–II) method for zircon from the Ai For-mation (Ronkin and Lepikhina, 2008) is inconsistentwith the series of isotopic-geochronological dates char-acterizing the rocks of the Berdyaush massif. As is ten-tatively mentioned in the quoted work, the obtainedresult presumably characterizes the impact of riftingduring the Mashak period on the isotopic systems in theminerals from the base of the Burzyan Group, but thedated minerals have not been described. We will notcomment on them until additional elucidation of theassumed phenomenon.

In concluding this part of the work, we would like toemphasize the following. Some researchers argued thatthe recent U–Pb (SHRIMP–II) dating of zircons fromboundary horizons between the Burzyan and Yurmatagroups suggest the position of the Lower–Middle Riph-ean boundary in the Uralian stratotype within the time

span of 1385–1370 Ma. According to StratigraphicCode (1992), any boundary of chronostratigraphic sub-division is the isochron plane, whose position can befixed by a deliberately selected stratotype point, and thetime-range of a chronostratigraphic unit is defined bythe succession of lower boundaries for stratigraphicsubdivisions of identical rank. Consequently, the U–Pb(SHRIMP–II) dates could be interpreted as an argu-ment in favor of positioning the Lower–Middle Riph-ean boundary at the level of 1370 ± 16 Ma according tothe results of the dating for the Mashak volcanicsbeginning in the Middle Riphean succession.

In our opinion, however, the U–Pb (SHRIMP–II)dates characterizing the Mashak zircons are insuffi-ciently reliable and precise (in particular, because of thelimited number of zircon samples whose U, Th, and Pbisotopic composition was analyzed in single points) tobe used as valid grounds for revising the isotopic age ofthe boundary that separates two Riphean erathems.Such a revision would be reasonable only after thecomprehensive study of the U–Th–Pb geochronome-ters from boundary horizons using the classical U–Pbmethod. At present, the isotopic age of the Lower–Mid-dle Riphean boundary should be retained at the tradi-tional level of 1350 ± 50 Ma.

PALEONTOLOGICAL CHARACTERIZATION OF THE BURZYAN GROUP

Organic remains represented by stromatolites andorganic-walled and silicified microfossils are wide-spread in the Lower Riphean deposits of Russia. Theyimpart a unique biostratigraphic character to the groupbut cannot be used for its subdivision because of thefacies factor influence.

An impoverished assemblage of organic-walledacritarchs discovered in the Lower Ai Subformationrepresents the oldest microfossils of the Burzyan Groupin the Urals. It includes some long-ranging forms andthe Riphean taxon Satka favosa (Kozlov et al., 1989;Yankauskas et al., 1989). A more diverse microbiota isdescribed from the Lower Kusa Subformation of theSatka Formation (Stratotype …, 1982; Veis et al., 1990;Sergeev and Lee Seong-Joo, 2004; Sergeev, 2006).Remains of long-ranging filamentous and coccoidalcyanobacteria (Gloeodiniopsis, Eosynechococcus,Palaeolyngbya, Eomicrocoleus, and others) are associ-ated in this assemblage with newcomers, namely withsphaeromorphic acritarchs Satka favosa, Kildinellahyperboreica, Leiosphaeridia crassa, L. bicrura,L. atava, Nucellosphaeridium minitum, Protosphaerid-ium densum, Leiominuscila minuta, Eomarginata stri-ata, Coniunctiophycus, Germinosphaera, Pterosper-mopsimorpha, (?)Granomarginata, Myxococcoidesminor, M. inornata, less common Palaeoanacistis vul-garis, Leiotrichoides, and Germinosphaera tadasii.The assemblage is barren of forms with spines and pro-cesses except for the problematic Micrhystridium sp.(Stratotype …, 1982), whose spines are most likely of



secondary origin (Sergeev, 2006). Information that thismicrobiota includes Gunflintia (Krylov and Sergeev,1986) has not been confirmed. Some representatives ofthe listed taxa occur in the Upper Kusa Subformation aswell.

The above microfossils from the Satka Formationwhose silicified varieties represent the Satka type ofmicrobiotas (Sergeev, 2006) are of a limited lateral dis-tribution. Beyond the Burzyan Group of the Taratashanticlinorium, they occur only in the Revet Subforma-tion of the Avzyan Formation, the terminal one in theMiddle Riphean stratotype of the Urals, and at the baseof the Svetlyi Formation, the middle one in the MiddleRiphean succession of the Uchur–Maya region ofsoutheastern Siberia (Veis and Vorob’eva, 2002; Ser-geev and Lee Seong-Joo, 2001). The diverse organic-walled microbiota confined to the Lower Riphean sub-tidal deposits in the mentioned region of Siberia, whichcontain the stromatolite assemblage of the Lower Riph-ean (Veis and Semikhatov, 1989), can be attributed tothe same type.

In the Bakal Formation of the Burzyan Group,organic-walled microfossils occur nearly in all silici-clastic members of the Lower- and Upper Bakal subfor-mations. In addition to many genera characteristic ofthe Satka Formation, those members yield coccoidalforms Myxococcoides minor, Leiominuscula minuta,along with the rarer Satka favosa and Eomarginata sp.These taxa occur in association with colonies of smallspheroids Palaeoanacistis vulgaris, remains of eukary-otic organisms Germinosphaera tadasii, species withmedian splits in their sheaths Leiosphaeridia bicrura,fragments of cyanobacteria Palaeopleurocapsa sp.,acritarchs with internal bodies Nucellosphaeridium andPterospermopsimorpha, rare sheaths Leiotrichoides,and ellipsoids Brevitrichoides.

Equivalents of the Burzyan Group recovered by drillholes in the Kama–Belaya aulacogen of the Cis-Uralsyield three successive microbiotas of organic-walledmicrofossils, which differ in composition from micro-biotas occurring in the type sections of the group(Stratigraphic …, 2000). Rare forms Leiosphaeridiajacutica ranging sometimes up to the Upper Ripheanhave been identified, in addition to long-ranging taxa,in equivalents of the Lower Ai Subformation. In thedesignated aulacogen, the diverse microbiota has beendiscovered in stratigraphic equivalents of the Satka For-mation upper part. In this microbiota, dominant long-ranging taxa are associated with typical Lower Ripheanforms Navifusa majaensis, Brevitrichoides bashkiricus,and Pellicularia ternata. The other associated taxa areLeiosphaeridia jacutica, Chuaria circularis, Recta cos-tata, Plicatidium latum, and Archaeoclada sp. alsoranging up to the Upper Riphean in some sections. TheLower Riphean Satka undosa, S. favosa, Brevitri-choides burzjanicus, and certain astratigraphic formsare known from equivalents of the Bakal Formationrecovered by the same boreholes. Veis and some other

researchers (Kozlov et al., 1998; Veis et al., 2001), whotook into consideration the large size of the designatedtaxa and the occurrence among them of microfossilsknown from the Upper Riphean Karatau Group, attrib-uted the respective part of succession in the Kama–Belaya aulacogen to the Upper Riphean. Nevertheless,the wide spectrum of correlation methods that havebeen used showed convincingly that the rock succes-sion under consideration is identical to the Burzyantype section and belongs to the Lower Riphean (Strati-graphic …, 2000; Kah et al., 2007).

The much wider lateral distribution is typical of theother Early Riphean microbiota from the shallow-watercarbonate facies that is termed the Anabar assemblageof silicified microfossils (Sergeev, 2006). Dominantforms of the assemblage are akinetes Archaeoellip-soides, associated short trichomes Filiconstrictosus,Orculiphycus, Partitiofilum, and coccoidal Myxococ-coides grandis. They occur in association with remainsof entophysalidacean and Synechococcus-like cyano-bacteria, chroococcacean cyanophytes Gloeodiniopsis,and empty sheath Siphonophycus, the taxa determiningcompositional specifics of microbiotas characteristic ofsecond half of the Early Proterozoic (Hofmann andSchopf, 1983). Some taxa of the assemblage range upto the Upper Riphean and Vendian, but in general themicrobiota in question dominates only in the Lowerand initial Middle Riphean deposits. In Russia, themost representative Anabar assemblage of microfossilsis known from the Lower Riphean Kotuikan Formationof the Anabar massif. Chronostratigraphic range of theformation is constrained by the dates of 1483 ± 3 Ma(the Rb–Sr age of unaltered glauconite from the under-lying deposits of the Ust-Ilya Formation; Gorokhovet al., 1991) and 1384 ± 2 Ma (the U–Pb age of badde-leyite from cross-cutting dyke; Ernst et al., 2000). Thewide distribution of the Anabar-type microbiotas in theEarly and initial Middle Riphean was not a result of therespective microorganisms’ evolution, but a conse-quence of specific global environment that appeared atthat time (Sergeev, 2006). Local development of theSatka-type microbiotas against this background waslikely controlled by the particular combination offacies-ecologic factors that is unknown as yet.

Finally, one more type of the Lower Riphean micro-biotas occurs in the Ust-Ilya Formation of the Anabarmassif, unaltered glauconite, from which has the Rb–Srdate of 1483 ± 3 Ma (Gorokhov et al., 1991). Micro-biota of this type is represented by organic-walledmicrofossils, most common of which are morpholog-ically simple filamentous and coccoidal formsLeiosphaeridia, Ostiana, Sphaerocongregus, Eosynecho-coccus, Leiosphaeroides, Eomarginata, Siphonophycus,Rectia, akinetes Brevitrichoides, and wide trichomesBotuobia. These forms are accompanied by large Chua-ria, Coniunctiophycus, Caudosphaera, Elatera, Eoso-lena, Plicatidium, Rugosoopsis, Aimia, Arctacellularia,Trachytrichoides, and Majasphaeridium of compli-cated morphology, and by thalluses Majaphyton and

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Ulophyton, all unknown from the preRiphean deposits(Veis and Vorob’eva, 1992; Veis et al., 2001). Outsidethe Anabar massif, nearly all microfossils listed aboveand some other “advanced” forms are widespread in theUchur–Maya region, being confined here to the middlepart of the Middle Riphean succession (Totta Forma-tion), and some of them range in this region up to theLakhanda Group of the Upper Riphean (Veis andVorob’eva, 1992; Veis et al., 2001).

The considered microfossils from the BurzyanGroup and its equivalents exemplify the unique bios-tratigraphic characteristic of the Lower Riphean depos-its that makes them different from the underlyingdeposits of the Lower Proterozoic. On the other hand,the spatial-geochronological distribution of concretetaxa representing the Lower Riphean microbiotasdepended to a great extent on facies-ecologic factorsand cannot be used as substantiation for the generalsubdivision of their host deposits.

Stromatolites widespread in most carbonate depos-its of the Burzyan Group contribute much to biostrati-graphic specifics of the Lower Riphean Erathem. Theoldest Burzyan Group stromatolites were responsiblefor the origin of successive bioherms occurring in theLower Kusa Subformation of the Satka Formation andrepresented by the Kussiella kussiensis buildups. Thistaxon, an important one among the Lower Ripheanstromatolites (Krylov, 1975; Stratotype …, 1982), wasformerly considered to be confined exclusively to theBurzyan level. Later on, it was established, however,that K. kussiensis occurs in the Uchur–Maya regionwithin the basal part of the Middle Riphean, being asso-ciated at this level with characteristic stromatolite mor-photypes, whereas in Australia this taxon was foundnear the top of the pre-Riphean deposits (Semikhatovand Serebryakov, 1983; Walter et al., 1988). In theLower Kusa Subformation of the Uralian sections,K. kussiensis is accompanied by less abundant formGongylina differenciata, and both taxa (the former onealways being dominant) occur in the Upper Kusa sub-formation in association with some endemic forms.

Some changes in compositions of stromatoliteassemblages under the influence of a changed sedimen-tation environment are recorded in the Lower SatkaSubformation. Former subtidal taxa became locallyassociated in respective periods with comparativelydeep-water morphotypes that populated settings belowthe storm wave base (Petrov and Semikhatov, 2005).These are representatives of the genus Conophyton(C. garganicus and endemic species of the genus) andstratiform buildups of the morphologically complicatedThyssagetacea subgroup (Vlasov, 1977). In the upperpart of the Burzyan Group (Berezovaya, Shuida, andGaevskii members of the Bakal Formation), deep-waterforms became universally dominant among stromato-lites. The Conophyton cylindricus prevailing at thislevel is accompanied by less abundant C. garganicus,C. lituus, Jacutophyton sp. and some local forms. (Kry-

lov, 1963, 1975; Komar et al., 1965; Stratotype …,1982). In the Berezovaya and Shuida members, theextensive Conophyton bioherms are 1.5 to 25-m thick,whereas large bioherms and subordinate biostromes fillin completely the 80-m-thick interval of the GaevskiiMember. In contrast, the uppermost strata of the BakalFormation contain only small Gongylina differenciatabuildups, characteristic of the Lower Riphean shallow-water deposits in Siberia (Semikhatov and Serebrya-kov, 1983). The Conophyton and associated Jacutophy-ton stromatolite forms are characteristic of the BurzyanGroup and its equivalents appear in the composite geo-logical record beginning since the lower part of theLower Riphean and occur higher throughout the Mid-dle and Upper Riphean, where they are associated withnew characteristic stromatolite morphotypes (Komaret al., 1965; Semikhatov and Raaben, 1994, 1996). Theappearance levels of Conophyton buildups in theBurzyan Group succession characterize the stages ofthe deepening of the basin.

Stromatolite assemblages from the Lower Ripheansubtidal deposits of the Uchur Group in the Uchur–Maya region are of different compositions (Semikhatovand Serebryakov, 1983; Serebryakov, 1975). Here, thewidespread taxa are Omachtenia omachtensis, Panis-collenia omachta, Stratifera omachtella, Gongylinadifferenciata, and Nucleella figurata occurring in asso-ciation with the rare Colonnella sp., single Kussiellakussiensis, and some endemics. The Uchur Group con-taining these stromatolites overlies heterogeneous rockcomplexes: the crystalline basement of theDzhugdzhur–Stanovoi region in the Siberian craton,the volcano-plutonic complex whose youngest rocksyield the U–Pb zircon ages of 1718 ± 1 to 1704 ± 5 Ma(Larin et al., 1997; Larin, oral communication), andlocal areas of the Uyan Group (700 to 1700 m thick)that has not been dated so far. The K–Ar ages of glauc-onites that have not been studied mineralogically corre-spond to 1520 and 1450 Ma in the lower part of theUchur Group, and to 1360 Ma in the upper one. Theunconformable upper boundary of the Uchur Groupseparates it from the overlying Aimchan Group of theMiddle Riphean. The local terminal subdivision of thelatter (Svetlyi Formation) yields silicified microfossilsof the Satka type (Sergeev and Lee Seong-Joo, 2001)and stromatolites that are attributed to indicative inter-regional taxa of the Middle Riphean (Svetliella svetlica,Sv. venusta, Baicalia aborigena, and B. inventa). TheK–Ar age of glauconite from the lower (Talyn) formationof the Aimchan Group corresponds to 1230–1210 Ma.

Some stromatolites known from the Uchur Groupoccur also in the Lower Riphean deposits of deeper sed-imentation settings on the Olenek Uplift and in theiranalogs in the Anabar massif (Komar, 1966; Semi-khatov and Serebryakov, 1983; Serebryakov, 1975). Inthe lower part of the respective succession that accumu-lated in the open shelf zone (the Ust-Ilya Formation ofthe Anabar massif and the Lower Kyutingda Subforma-tion of the Olenek Uplift), Kussiella kussiensis and



Colonnella discreta accompanied sometimes by Cono-phyton garganicus, local Jacutophyton species, andNucliella fibrosa on the Olenek Uplift were found,whereas in the Anabar massif they are associated withConophyton cylindricus and rare endemic species ofgenera Kussiella and Colonella. Upper subtidal depos-its of the succession under consideration (the UpperKyutingda Subformation and its equivalents in the Ana-bar massif) are dominated by stromatolites of the strat-iform type (Stratifera undata, Gongylina differenciata,and Nucliella figurate). In the Anabar massif, theseforms occur in association with Kussiella kussiensis.The Rb–Sr and K–Ar dates of 1483 ± 3 Ma and 1459 ±10 Ma, respectively, are known for unaltered glauconitefrom the Ust-Ilya Formation (Gorokhov et al., 1991),whereas baddeleyite from dyke crosscutting the abovedeposits is dated at 1384 ± 2 Ma by the U–Pb method(Ernst et al., 2000).

In evaluating the biostratigraphic significance ofstromatolites from the Burzyan Group, one should takeinto account the fact that they differ from the antedatingLower Proterozoic assemblages not only in the compo-sition of their species, but also in the lateral distributiontrends. The distribution areas of the preRiphean stro-matolite assemblages are limited as a rule, and domi-nants of the assemblages are endemic species and mostgenera, which do not occur in the Riphean strata. In thesubordinate group of the preRiphean genera, whichcross the Riphean base and occur in higher strata(Conophyton, Jacutophyton, Stratifera, Paniscollenia,Colonnella, and others), the long-ranging species rep-resent an exception (Semikhatov, 1978; Semikhatov,1991; Semikhatov and Raaben, 1994, 1996). Beingimportant for biostratigraphic characterization of theLower Riphean as a whole, stromatolites are inappro-priate, however, for subdivision of this erathem intolower-rank subdivisions. As is shown above, changes inthe taxonomic composition of their assemblages,observable upward in the sections, depended to a greatextent on facies and ecological factors, thus beingdeprived of general stratigraphic significance.

In a series of published works, the appearance ofmicrofossils Satka favosa, Leiosphaeridia crassa,L. bicrura, Myxococcoides grandis, Pterosphermopsi-morpha, (?) Granomarginata, Eomarginata, Archae-oellipsoides, and some others, which occur in the Bakaland Satka formation of the Southern Urals (single spec-imens of Satka favosa are known also from the Ai For-mation), has been regarded as a promising biostrati-graphic substantiation of the Burzyanian Erathem base.However, these microfossils appear away from thelithostratigraphic base of the Burzyan Group and alongwith biofacies of the other microbiota types that preventsusing them as markers of the Lower Riphean base.

As a palliative for solving the problem, one can usedata on the Siberian hypostratotype of the Riphean, i.e.,on the succession of the Uchur–Maya region (Semi-khatov and Serebryakov, 1983). The event boundary

separating the Uchur Group from the underlying com-plexes (gneisses of the Siberian platform basement,volcano-plutonic belt that terminated its evolution1718 ± 1 to 1704 ± 5 Ma ago, and the Uyan Group oflocally occurring volcanogenic-siliciclastic deposits) isof great significance for the traditional substantiation ofthe Riphean base in the Uchur–Maya region. Stroma-tolitic dolomites deposited 60 to 70 m above theBurzyan Group base contain the shallow-water assem-blage of interregional stromatolite taxa Omachteniaomachtenis, Stratifera omachtella, more rare Paniscol-lenia omachta, single Gongylina differenciata andKussiella kussiensis, which are known from the LowerRiphean deposits of Siberia, but do not occur in the pre-Riphean strata, as is shown above. It is reasonable tokeep in mind that mineralogically unstudied glauco-nites from the Gonam Formation base are dated by theK–Ar method at 1520 and 1460 Ma, whereas in theoverlying Omakhta Formation their age corresponds to1360 Ma. It is likely therefore that after investigation ofmicrofossils between the occurrence level of the abovestromatolites and the base of the Uchur Group, the sed-imentary succession exposed on the right side of theUchur River valley near the Suklan Creek mouth couldbe regarded as appropriate for fixing the stratotype ofthe lower boundary of the Lower Riphean (Burzyanian)Erathem.

C-O AND Sr-ISOTOPIC SYSTEMATICS OF CARBONATE ROCKS

FROM THE BURZYAN GROUP

Carbonate rocks (especially limestones) are a basicsource of the Rb–Sr and C–O chemostratigraphic infor-mation, as they retain until the present time, in certainconditions, the initial isotopic characteristics of Rb–Srand C–O isotopic parameters in the sedimentation envi-ronment. The main factor that changed these character-istics was the influence of the low-temperature diage-netic and epigenetic fluids originating during the trans-formation of aluminosilicates and dispersed organicmatter in associated siliciclastic and clayey carbonatesediments. As a result, limestones and dolomitesbecome enriched in Mn, Fe, Pb, Th, radiogenic 87Sr, butdepleted in 13C, 18O and U, relative to concentrations ofthese components in marine carbonate sediments(Drever, 1982; Chaudhuri and Clauer, 1993; Knollet al., 1995). Noticeable changes in the isotopic compo-sition of some elements (primarily of carbon and oxy-gen) originate under the influence of underground andmeteoric waters, which also differ in composition fromseawater and penetrate into carbonate rocks during thetectonic upwarping of the study region. Consequently,in carbonate generations of limestones and dolomites,their own isotopic and geochemical characteristics areshifted toward those that existed in epigenetic fluids.Thus, the Mn/Sr, Fe/Sr, δ18é values and covariation ofthese parameters, on the one hand, and between themand 87Sr/86Sr and δ13ë values, on the other, can be used

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for the selection of the least altered (“best”) samples ofcarbonates (Brand and Veizer, 1980; Banner and Han-son, 1990; Gorokhov et al., 1995; Semikhatov et al.,2004; Kuznetsov et al., 2006). Selecting samples forchemostratigraphic and isotopic-geochronologicalresearch during this study, we gave preference to thoselimestones of the succession, which lack visible indica-tions of secondary alteration, have the lowermost SApercentage, and are not associated with siliciclasticbeds and are sufficiently remote from the erosionboundaries.

C-O Isotopic Systematics

The carbon concentration, high in limestones anddolomites, but low in epigenetic fluids, favors retainingthe δ13ë initial values in rocks during epigenesis. Nev-ertheless, the organic diagenesis, i.e., the interactionbetween carbonates and fluids, which are enriched inthe light carbon isotope 12ë released during the decom-position of dispersed organic matter, the product of thelife activity of photosynthetic microorganisms, couldcertainly be responsible for secondary changes of δ13ëin carbonates (Veizer et al., 1992; Knoll et al., 1995;Knoll and Semikhatov, 1998). From this point of view,it is important that shales of the Satka and Bakal forma-tions contain on average 0.4 to 1.5% of dispersed fineorganic matter (Rykus et al., 1993; Krupenin et al.,1993), whose decomposition could enrich the epige-netic fluids in isotope 12C. Sedimentation in deep stag-nant water is most favorable for burial of that matter inthe fine-grained siliciclastic sediments.

Criteria that point to retentivity of the C-isotopic ini-tial systems in carbonates are the ordinary valuesMn/Sr < 10 and/or δ18O > –10‰ (Knoll et al., 1995;Kaufman and Knoll, 1995). On the other hand, δ13ë incarbonates not satisfying these criteria can be insignif-icantly shifted toward positive or negative values rela-tive to this parameter in the “best” samples from thesame formations. It is very likely therefore that underlow-temperature conditions the epigenetic fluids areable to extract Mn and Fe from the siliciclastic admix-ture without any perceptible fractionation of the carbonisotope composition. A process of this kind was sug-gested for clay limestones of the Upper Riphean KatavFormation in the Southern Urals (Kuznetsov et al.,2006).

The δ18é value in unaltered (“best”) marine carbon-ates of the Upper Proterozoic corresponds to –6.5 ±2.5‰ V-PDB (Veizer and Hoefs, 1976; Ray et al.,2003), being below –10‰ in rocks that experiencedepigenetic alterations (Kaufman and Knoll, 1995).Researchers commonly use exactly this threshold valueof δ18é when separating the “best” and altered samplesof Precambrian limestones. Analyzing C-isotopic sys-tems of carbonates from the Burzyan Group, we used asbefore the stricter geochemical criteria to select the“best” samples: Mn/Sr < 4, Fe/Sr < 10, δ18é > –10‰for limestones and Mn/Sr < 6, Fe/Sr < 15, δ18é > –10‰

for dolomites (Semikhatov et al., 2004; Kuznetsovet al., 2006).

During this study, C- and O-isotopic compositionwas analyzed in the VSEGINGEO on the modernizedmass spectrometer MI-1201B in a single ëé2 volume,obtained after the dissolution of samples approximately20 mg in weight in concentrated orthophosphoric acidunder temperature 25°ë during 2 (limestone samples)and 72 hours (dolomite samples). Cleaning of ëé2 wascarried out in accordance with the conventional proce-dure (Hostenberger and Herman, 1984). Working stan-dard MCA-7 attested in the TSNIGRI and VIMS-7(δ13ë = +2.2‰, δ18é = –8.8‰)) was calibrated versusthe V-PDB scale, using the standards NBS-18, KH-2and TKL. Analytical uncertainty was not greater than±0.2‰ (1σ) for δ13ë and ±0.4‰ (1σ) for δ18é. The C-and O-isotopic composition was studied in 57 samples,nearly half of which are from the carbonate Satka For-mation, with the thickest one in the Burzyan Group.One third of the studied samples are from the BakalFormation, and about 17% of samples characterize car-bonates of the Suran Formation.

Satka Formation. C- and O-isotopic characteristicsare studied in 26 samples of the formation’s carbonates,19 of which (11 dolomites and 8 limestones) have a lowSA content (3.3% in average) are from the Upper SatkaSubformation. In addition, we analyzed 7 samples fromthe Lower Kusa Subformation and adjacent part of theUpper Kusa Subformation (Fig. 2, Table 2), where theSA content in rocks is higher. Rocks from the formationmiddle part with high SA and Fe content have not beenanalyzed.

Among carbonate rocks of the Satka Formation,limestones of the Kazymovskaya Member are of partic-ular importance. They contain 0.8 to 4.8% SA, and theirMn/Sr and Fe/Sr ratios are lower than the threshold val-ues (0.01–0.03 and 0.1–1.0 respectively), whereas δ18évalues ranging from –8.6 to –6.3‰ are close to theparameters characterizing the unaltered carbonate sed-iments of the Proterozoic (Fig. 3a). In limestones of themember, δ13ë values range from –0.6 to 0.0‰ andapparently characterize the carbon isotope compositionin the sedimentation environment. In contrast, the Fe/Srratio ranging from 14.8 to 31.5 is characteristic of claylimestones from the Lower and Upper Kusa subforma-tions. The δ18é values in these rocks vary from –11.3 to–10.6‰, although their Mn/Sr ratios are below the crit-ical level. The δ13ë of limestones from both subforma-tion changes from –0.9 to –0.4‰ but cannot beregarded, according to geochemical criteria, as charac-terizing this parameter in the Early Riphean seawater(Fig. 4).

In the dolomites of the Satka Formation, the Mn/Srratio is below the critical value only in one sample fromthe Lower Kusa Subformation, but its Fe/Sr ratio ismuch higher than the critical level. In three samplesfrom the formation with a low SA content (0.7–2.3%),the Fe/Sr ratio is close but still higher than the critical



level and ranges from 11.1 to 13.6. Simultaneously, theδ13ë values range from –0.9 to +0.1‰, while the δ18éof the dolomites varies from –9.6 to –3.4‰. The estab-lished difference between δ18é values in the calcite–dolomite pairs of the studied samples is perceptiblylower than Δδ18écal.-dol. characterizing the oxygen iso-tope’s fractionation under low temperatures in the equi-librium system of these two minerals (Fritz and Smith,1970). Consequently, the minerals of the studied dolo-mites are not cogenetic and have been formed mostlikely at the stage of sediment dolomitization under theinfluence of meteoric or other fluids enriched in 16O(Veizer and Hoefs, 1976; McKenzie, 1981).

Bakal Formation. C- and O-isotopic compositionis determined in 17 limestone samples from the Bere-zovaya and Gaevskii members, and in four dolomitesamples from the Shuida and Gaevskii members. Car-bonates of the Shikhan and Upper Bakal membersenriched in SA have not been analyzed. According togeochemical criteria, all limestone samples from theBerezovaya Member are of the “best” type, having aSA content from 0.2 to 0.7%, Mn/Sr = 0.01–0.13,Fe/Sr = 0.2–1.4, and δ18é from –8.1 to –6.0‰. Theδ13ë value in these rocks ranges from –0.3 to 0.0‰ thatalmost coincides with the variation range of this param-eter in the least altered limestones from the Kazy-movskaya Member of the Satka Formation. Allgeochemical parameters of two dolomite samples fromthe Shuida Member are higher than the critical values,

and δ13ë in these samples corresponds to –0.8‰ and–1.0‰. At the base and top of the overlying Gaevskiimember, there are 5- to 7-m-thick beds of dolomiteswith very high Mn/Sr and Fe/Sr ratios. In the lower bed,these ratios correspond to 160 and 2240, respectively,declining to 20.7 and 280 in the upper one. In the mid-dle part of the Gaevskii Member (about 65-m thick)that is composed of pure and dolomitic limestones(Mg/Ca 0.011–0.121), only one limestone sample fromthe lower part of the interval showed comparatively lowvalues of Mg/Ca = 0.073, Mn/Sr = 0.6, and Fe/Sr = 9.9,which exceed nevertheless their critical levels; δ13ë inthis sample is equal to –0.7‰. Close to siliciclasticmembers underlying and overlying the Gaevskii Mem-ber, the Mn/Sr and Fe/Sr ratios increase up to 1.3 and44.4, respectively, in the basal interval of the limestonesuccession above the aforementioned sample, and up to2.1 and 64.7 at the top. The δ13ë in limestones changestherewith from –0.8 to –2.1‰, being as low as –2.5‰ inthe dolomite bed with δ18é = –11.1‰ that crowns thelimestones. The considered data and correlationbetween δ18é and δ13ë in carbonates of the GaevskiiMember (Fig. 3b) clearly point to the epigenetic recrys-tallization of limestones under the influence of fluidsderived from underlying and overlying low-carbon-aceous shales.

The Suran Formation of the Yamantau anticlino-rium includes thick carbonate successions in the lower(Minyak) and upper (Lapyshta) subformations only.

Table 2. The Rb–Sr analytical data on carbonate fractions enriched in primary material from “best” limestones of the Satkaand Bakal formations

Sample no. Rb, ppm Sr, ppm 87Rb/86Sr (87Sr/86Sr)measured (87Sr/86Sr)initial

Bakal Formation (Berezovaya Member, Upper Bakal subformation)

UB-101 0.01 1040 0.0001 0.70474 0.70474

UB-121 0.05 960 0.0002 0.70479 0.70479

UB-198 0.15 1000 0.0004 0.70457 0.70456

UB-209 0.06 990 0.0002 0.70462 0.70462

UB-62 0.07 770 0.0003 0.70463 0.70463

UB-65 0.04 830 0.0001 0.70467 0.70467

UB-76 0.04 990 0.0001 0.70471 0.70471

UB-81 0.04 790 0.0002 0.70481 0.70481

Satka Formation (Kazymovskaya Member, Upper Satka Subformation)

UC-79 0.58 2450 0.0005 0.70473 0.70472

2-21 0.10 2740 0.0001 0.70465 0.70465

UC-78 0.03 2670 0.0001 0.70460 0.70460

2-18 0.12 2180 0.0002 0.70470 0.70470

2-16 0.12 2340 0.0002 0.70466 0.70466

UC-74 0.04 1675 0.0001 0.70468 0.70468

UC-73 0.54 1490 0.0011 0.70482 0.70480

Note: in calculation of the 87Sr/86Sr initial ratios, age of rocks from the Bakal and Satka formations was assumed to be 1550 and 1430 Ma,respectively.

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In the Minyak Subformation, C- and O- isotopic compo-sition has been studied in the basal 40-m-thick dolomite(Mg/Ca 0.625–0.645) member exposed in proximity tothe Suran village, where rocks of the member contain15.3 to 23.1% SA, 650 to 6280 ppm Fe, being somewhatdepleted in Mn (240–310 ppm), whereas very low δ18évalues range in the member from –14.6 to –11.6‰. Theoverlying limestones of the member (45 m thick), whichare exposed near the Ismakaevo village, have a lower SAcontent (4.9–16.6%) than the underlying dolomites,being close to the latter in Mn concentration but consid-erably enriched in Fe (1490–6010 ppm), whereas theδ18é values range in these rocks from –13.3 to –15.3‰.Carbonate generations with similarly low δ18é valuescan be formed in the course of the high-temperaturerecrystallization of rocks and oxygen isotope fraction-ation in a partially closed system, but δ13ë values do notchange significantly under such conditions (Veizer et al.,1992). That was apparently the reason why δ13ë valuesof –0.7 to +0.1‰ in the Suran limestones are close in thisparameter to the “best” limestones of the Satka andBakal formations. In the Minyak limestones, δ13ë iswithin the narrow variation range from –0.6 to –0.2‰and corresponds to +0.2‰ in the dolomites (Fig. 3).

Limestones of the Lapyshta Subformation, crown-ing the Suran Formation section in proximity to theKartalinskaya Zapan village, have very low Mg/Ca(0.003 to 0.41) and satisfactory Fe/Sr (3.0 to 4.5) ratios.The Mn/Sr ratio approaching from above the criticallevel ranges in these limestones from 0.24 to 0.41, butδ18é exceeds the threshold value and ranges from –14.6to –11.6‰. Thus, data on rocks of the Suran Formationadd nothing new to the inference about δ13ë variationsin the Burzyan ocean. As is established based on theresults obtained for the least altered (“best”) limestonesfrom the Kazymovskaya and Berezovaya members ofthe studied sedimentary succession, δ13ë varied in thatpaleobasin from –0.7 to 0.0‰ (Fig. 4). The insignifi-cant variation range of δ13ë values suggests that thepaleogeography and the paleogeodynamic regime ofthe paleobasin did not change significantly during theSatka and Bakal time. Judging from the considereddata, lowered δ13ë values in carbonate rocks from theLower and Upper Kusa subformations of the Satka For-mation and from the Shuida and Gaevskii members ofthe Bakal Formation (as low as –2.5 and–2.1‰, respec-tively) characterize changes in the C-isotopic systems

δ13C

–0.5

0

–1.0

–2.0

–3.0–16 –12 –8 –4 –2

δ18O

(a)LA

δ13C

–0.5

0

–1.0

–2.0

–3.0–16 –12 –8 –4 –2

δ18O

(b)LA

δ13C

–0.5

0

–1.0

–2.0

–3.0–16 –12 –8 –4 –2

δ18O

(c)LA

LA1 2 3 4 5

Fig. 3. Covariation of δ13ë and δ18é values in carbonaterock samples from the Satka (a), Bakal (b) and Suran (c)formation in comparison with field of the same parametersin the least altered marine carbonates of the initial Late Pro-terozoic (Early Riphean): (1) limestones satisfying geochem-ical criteria of retentivity Mn/Sr ≤ 4 and Fe/Sr ≤ 10; (2) lime-stones unsatisfying the same criteria; (3) dolomites satisfyinggeochemical criteria of retentivity Mn/Sr < 6 and Fe/Sr < 15;(4) dolomites unsatisfying the same criteria; (5) field of datapoints characterizing the least altered marine carbonates ofthe Lower Riphean.



of carbonates, which originated under the influence oforganic diagenesis and a temperature rise during tec-tonic downwarping.

Sr-isotopic Systematics

The Sr-isotopic systematics of carbonates from theBurzyan Group is studied in 53 rock samples: 25 fromthe Satka, 20 from the Bakal, and 8 from the Suran for-mations (Fig. 4). As in the case of C-isotopic systemat-ics, the least altered samples suitable for assessment ofthe 87Sr/86Sr ratio in the sedimentation environment andchemostratigraphic considerations have been selectedusing such geochemical criteria as the Mn/Sr, Fe/Sr,Ca/Sr and Mg/Ca ratios (Brand and Veizer, 1980; Gor-okhov et al., 1995). Because of considerably differentSr and C concentrations in epigenetic fluids, the Rb–Srisotopic systems in carbonates are less stable againstthe postsedimentary reworking than the C-isotopic sys-tems (Banner and Hanson, 1990), and consequently thethreshold Mn/Sr and Fe/Sr values should be consider-ably lower in the case of the Rb–Sr isotopic study thanduring the C-isotopic systematics. According to ourexperience in studying the Precambrian carbonaterocks, strict geochemical criteria useful for selectingthe best limestone samples for the Rb–Sr isotopic anal-ysis are as follows: Mn/Sr < 0.2, Fe/Sr < 5.0, Mg/Ca <0.024 and δ18O > –10.0‰ (Gorokhov et al., 1995; Kuz-netsov et al., 1997, 2003·, 2006; Semikhatov et al.,2002). In the case of dolomites, we determined theabove criteria individually for each lithostratigraphicunit and only then, when the Mn/Sr and Fe/Sr ratiosreveal positive correlation with the 87Sr/86Sr ratio. Sucha correlation is not established for dolomites of theBurzyan Group (Kuznetsov et al., 2008, Fig. 3).

After preliminary treatment of limestone samplesfrom the Burzyan Group in 1N ammonium acetatesolution, we dissolved the residue in the 10% aceticacid in order to study the Rb–Sr systematics of carbon-ate fraction. The preleaching in ammonium acetate is anecessary procedure, as it partly removes the epigeneticcarbonate generations represented mostly by magne-sian calcite or low-magnesian dolomite, which areenriched in Rb and radiogenic 87Sr (Gorokhov et al.,1995; Kuznetsov et al., 1997, 2003b, 2008). Despitepossible epigenetic recrystallization, the carbonatefraction dissolved in acetic acid at the second stage isperceptibly enriched in primary carbonate material.The Rb concentration in the enriched carbonate frac-tion of the Burzyan limestones and dolomites is onaverage 10 and 28 times lower, respectively, than in epi-genetic carbonate generations (Kuznetsov et al., 2005,2008). Difference between the 87Sr/86Sr ratio measuredin epigenetic carbonates and fraction enriched in pri-mary carbonate material changes from 0.0001 to0.0104 in limestones and from 0.0014 to 0.0199 indolomites. The Sr isotopic ratios of limestones anddolomites reveal a positive correlation with the Mn/Sr(r = 78, n = 20) and Fe/Sr (r = 68, n = 20) ratios that evi-

dences the noncogenetic origin of carbonate genera-tions and points to the enrichment of secondary gener-ation in Mn, Fe and radiogenic 87Sr in the course of epi-genetic recrystallization. The 87Sr/86Sr values quoted inthis work characterize carbonate fractions of lime-stones and dolomites enriched in primary carbonatematerial.

The isotope dilution method with mixed spike87Rb + 84Sr is used to determine the Rb and Sr concen-trations in the studied samples. The Sr isotopic compo-sition is measured on the Triton TI multicollector massspectrometer in the static acquisition mode. During theanalytical period, average 87Sr/86Sr ratio normalized to86Sr/88Sr = 0.1194 corresponded in standard samplesNIST SRM 987 and EN-1 to 0.710262 ± 0.000005(2σaver., n = 34) and 0.709191 ± 0.000005 (2σaver., n =17) respectively.

Satka and Bakal formations. In the BurzyanGroup, the least altered carbonates characterizing theSr-isotopic parameter of sedimentation environmentoccur only in the Kazymovskaya Member of the Satkaformation and in the Berezovaya Member of the BakalFormation (Table 2). Limestones of both members sat-isfy the strict geochemical criteria of the isotopic sys-tem retentivity, have low SA content (0.9–3.5% in theKazymovskaya Member and 0.2–0.7% in the Bere-zovaya member), and their carbonate fractions havebeen analyzed after preliminary leaching in 1N solutionof ammonium acetate. In the enriched carbonate frac-tion of the “best” limestone samples from the members,the Rb concentration is not higher that 0.15 ppm as arule, being as high as 0.54 and 0.58 ppm in two casesonly (Table 2). Because of the high Sr concentration inthe analyzed samples (770–2740 ppm), the incrementof the 87Sr/86Sr ratio in response to the 87Rb radiogenicdecay in limestones of the Satka and Bakal formationsis insignificant. The increment corresponds to 0.00002in one case and 0.00001 in two others, being close tozero in all other analyzed fractions. The 87Sr/86Sr initialratio in the “best” samples from two stratigraphicallydisjoint members of the Burzyan Group is practicallyidentical and ranges from 0.70460 to 0.70480 in theKazymovskaya Member (1550 ± 30 Ma) of the SatkaFormation and from 0.70456 to 0.70481 in the Bere-zovaya Member (1430 ± 30 Ma) of the Bakal Formation(Kuznetsov et al., 2005, 2008). The above values arewell consistent with the 87Sr/86Sr ratios known for theleast altered carbonates of the Lower Riphean from theother regions of the world (Pokrovskii and Vinogradov,1991; Gorokhov et al., 1995; Hall and Veizer, 1996;Ray et al., 2003; Kah et al., 2007), where they arewithin the range of 0.7046–0.7050 and partially charac-terize samples analyzed without preliminary treatmentin ammonium acetate.

The Rb–Sr systems of clay limestones and dolo-mites from the Satka and Bakal formations, which con-tain perceptible admixture of siliciclastic materialsometimes, are apparently modified because of the

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SEMIKHATOV et al.

Bur

zyan

Yur

mG

roup

Ai

Satk

aB

akal

Zg

Z-K

Form

atio

n

St1

St2

St3

St5

Low

erU

pper

Subf

orm

atio

n

St5

-1S

t5-2

St5

-3M

akar

ovo

109876

5

4

3

2

1

Sr4

Form

atio

n

Subf

orm

atio

n

Zg

Z-K

Taratashanticlinorium

Yamantauanticlinorium

Bol

.In.

Sura

nY

usha

Mas

hak

1

2

3

4

5

Sr1

Sr2

Sr3

Sr4

Sr5

Mem

ber

500 m

δ13C

–2 –1 0 +1

0.70456–0.70481

0.70460–0.70480

δ13C

–2 –2

–2

–1 –1

–1

+1 +1

+1

0 0

0

0.70460

0



radiogenic 87Sr interchange between the carbonate frac-tions and siliciclastic material of the rocks and/or asso-ciated siltstones and shales. This is evident from thegeochemical criteria characterizing retentivity of isoto-pic systems in samples from the Satka Formation(Table 1). As a result of the above interchange, the87Sr/86Sr ratio ranges correspond to 0.71236–0.72932 inthe Lower Kusa Subformation, to 0.70972–0.71933 inthe Upper Kusa Subformation, and to 0.70616–0.73042in the Upper Satka Subformation (Kuznetsov et al.,2008). Samples from these subformations reveal highFe/Sr (up to 178) and Mn/Sr (up to 7.5) ratios alongwith lowered (down to –11.3‰) δ18é values. The Rb–Sr systems of limestones and dolomites from the Gae-vskii and Shuida members of the Bakal Formation havebeen disturbed by process of metasomatic sideritizationunder the influence of the Fe-bearing elision fluids dur-ing burial of sediments. The 87Sr/86Sr ratio in alteredepigenetic limestones and dolomites of the Bakal For-mation varies broadly from 0.70732 to 0.72923; indolomites most intensively reworked near the sideritebodies, it varies from 0.73765 to 0.74168 (Kuznetsovet al., 2005).

The Suran Formation. In all studied limestonesfrom this formation, the Mn/Sr varies from 0.23 to 1.92,being greater in general than the threshold value; the Fe/Srratio changes from acceptable (2.1) to high (33.0) values,and parameter δ18é ranges from –15.3 to –11.6‰, greatlyexceeding the threshold level. According to all thesecharacteristics, the rocks are considerably altered, andthe Suran Formation is lacking limestones characteriz-ing the 87Sr/86Sr ratio in the sedimentation environment.Variation ranges of the 87Sr/86Sr ratio correspond to0.70584–0.71090 in limestones of the Minyak Subfor-mation and to 0.70625–0.70843 in carbonates of theLapyshta Subformation. In dolomite samples from theformer subformation with low δ18é value (–13.9‰),the 87Sr/86Sr ratio is as high as 0.73805 (Kuznetsovet al., 2008).

CONCLUSIONS

The Burzyan Group representing type succession ofthe Lower Riphean in the Southern Urals is bounded atthe base and top by significant historical-geologicboundaries. The lower boundary separates the Taratashcrystalline complex whose rocks underwent the termi-nal metamorphic stage of amphibolite facies 1800.8 ±2.6 Ma ago (Sindern et al., 2005, 2006) from rudaceoussedimentary rocks of the Ai Formation hosting basaltsin its middle part, which contain zircons having the U–Pb date of 1615 ± 45 Ma (Krasnobaev et al., 1992). Theisotopic age of this boundary is estimated at the level of

1650 ± 50 Ma based on the results of telecorrelation(Semikhatov et al., 1991; Krasnobaev et al., 1992;Stratigraphic …, 2000). In the Taratash anticlinorium,the upper boundary of the Burzyan group separates theBakal Formation from quartzites of the Middle RipheanZigalga Formation that discordantly rests on the under-lying strata. In the Yamantau anticlinorium, the upperboundary of the Burzyan Group coincides with theunconformity between siliciclastic sediments of theYusha Formation and volcanogenic-sedimentary rocksof the Mashak Formation that is basal one in the MiddleRiphean type succession and contains zircons dated bythe U–Pb (SHRIMP-II) method at 1370 ± 16 Ma. In thiswork, we retain the upper boundary of the group at thetraditional level of 1350 ± 50 Ma.

The C-isotopic characteristics of the least alteredcarbonate samples from the group show that magnitudeof δ13C variations in different areas of the Burzyanpaleobasin was low, corresponding to –0.4 ± 0.3‰. Theδ13C values in seawater varied from –0.6 to 0.0‰ inthe terminal Satka time (ca. 1550 ± 30 Ma ago) andfrom –0.7 to 0.0‰ in the mid-Bakal time (ca. 1430 ±30 Ma ago).

In the opinion of some researchers, the low δ13C val-ues in the Early Riphean seawater characterize thelong-lasting “global stasis” of carbon isotope composi-tion in the World Ocean that commenced 2.06 Ga agoand terminated about 1.25 Ga ago (Brasier and Lindsay,1998; Lindsay and Brasier, 2002; Kah et al., 2001; Bar-tley et al., 2007). According to this hypothesis, the δ13ëvalue in the ocean was not greater than 0 ± 1‰ duringthe designated time span, and growth of this parameterup to +2‰ started only after 1.25 Ga. If the hypothesisof long-lasting global stasis of the carbonate carbonisotopic composition is correct, then C-isotopechemostratigraphy could be used as a reliable tool forthe recognition of the boundary between the Lower andMiddle Riphean deposits. However, the hypothesis isimproperly substantiated, based on data characterizinga limited number of regional successions and does nottake into consideration the C-isotopic characteristics ofthe Lower Riphean carbonate formations in the AnabarUplift of Siberia (Knoll et al., 1995), the Belt Super-group in Rocky Mountains of the United States (Halland Veizer, 1996), the McArthur Group in Australia(Veizer et al., 1992; Lindsay and Brasier, 2000), and theSemri Group in India (Ray et al., 2003). In the carbon-ate rocks of the stratigraphic subdivisions listed above,the δ13ë values broadly range from –7.1 to +3.1‰.Such a great variation range of the values is to a greatextent a consequence of epigenetic alterations in thecarbonates. In the least altered samples of the latter,which can be selected using the geochemical criteria

Fig. 4. C- and Sr-chemostratigraphic characteristics of carbonate rocks from the Burzyan Group: (1) limestones satisfying geochem-ical criteria of retentivity Mn/Sr < 4, Fe/Sr < 10; (2) limestones unsatisfying the same criteria; (3) dolomites satisfying geochemicalcriteria of retentivity Mn/Sr < 6, Fe/Sr < 15, δ18é > –10‰; (4) dolomites unsatisfying the same criteria; (5, 6) 87Sr/86Sr variationsin the “best” (Mn/Sr < 0.2 and Fe/Sr < 5.0) limestone samples (other symbols and abbreviations as in Fig. 2).

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SEMIKHATOV et al.

applied in this work to carbonates of the BurzyanGroup, δ13ë values are more than two times lower andrange from –1.8 to +3.1‰ (Fig. 5).

It is also important that the aforementioned LowerRiphean successions include intervals of carbonaterocks with persistently positive or negative δ13ë values.The negative values (–1.8 to 0‰, Ray et al., 2003) aretypical of the Rohtas Formation in the Semri Group ofcentral India, where the U–Pb age of zircons from tuffhorizons of the formation corresponds to 1602 ± 10 and1599 ± 8 Ma (Rasmussen et al., 2002). Carbonates withnegative δ13ë values (–1.6 to –0.7‰, Knoll et al., 1995)are characteristic as well for the Kotuikan Formation inthe Anabar Uplift, where its age limits are constrainedby the dates of 1483 ± 5 Ma (Gorokhov et al., 1991) and1384 ± 2 Ma (Ernst et al., 2000).

Persistently positive δ13ë values are typical of theNimbekhera (+2.0 to +3.1‰) and Kairahat(+0.5…+2.7‰) formations (Ray et al., 2003) sand-wiched in the Semri Group between metamorphicrocks 1700 Ma old and tuff horizons having the U–Pbzircon ages of 1631 ± 5 and 1628 ± 8 Ma (Rasmussen etal., 2002). Steadily positive δ13ë are established as wellin carbonates of the Newland (+0.4 to +2.5‰), Oltin(+1.3 to +1.7‰) and Mount Shields (+0.3 to +1.3‰)

formations, subdivisions of the Belt Supergroup inRocky Mountains of the United States (Hall and Veizer,1996). The supergroup includes synsedimentary sills inits basal and terminal horizons, which are dated at1468 ± 3 and 1401 ± 6 Ma, respectively (Sears et al.,1998; Ross and Villeneuve, 2003). Negative to positiveδ13ë values are recorded in carbonates of the Helena andSiyeh formations of the same supergroup (from –0.8 to2.5‰; Hall and Veizer, 1996) and of the McArthurGroup in Australia (from –1.6 to 2.2‰ δ13ë; Veizeret al., 1992; Lindsay and Brasier, 2000) representingthe Lower Proterozoic–Lower Riphean boundarydeposits dated at 1700–1600 Ma (Page et al., 2001).

Using our geochemical criteria of the C-isotopicsystem retentivity in the Precambrian carbonates, wecan state therefore that amplitude of the δ13ë variationsin the Early Riphean marine sediments was up to 4–5‰, i.e., comparable with that of the Middle Ripheantime (Fig. 5). Consequently, it is necessary to revise thepostulated time span of the global stasis in the carbon-ate carbon isotopic composition that postdated the larg-est Jatulian (Lomagundi) excursion of positive δ13ëvalues and started 2.06 Ga ago and terminated prior to1.28 Ga ago (Brasier and Lindsay, 1998; Lindsay andBrasier, 2002; Kah et al., 2001; Bartley et al., 2001).Data on the Lower Riphean successions in Siberia,

δ13C

+4

+2

0

–2

–4900 1000 1200 1400 1600 1700

Ma

Riphean

Late Middle EarlyPR1

1110

9

8

7

6

5

4

3

2

1

21

Fig. 5. Variations of δ13ë in carbonate deposits of the Lower Riphean: (1) limestones with Mn/Sr < 10 δ18é > –10‰; (2) dolomiteswith Mn/Sr < 10 and δ18é > –10‰.Numbers in the figure denote the following lithostratigraphic subdivisions: (1) McArthur Group, Australia (Lindsay and Brasier,2000); (2) Semri Group, Central India (Ray et al., 2003); (3) Satka Formation, Southern Urals (this work); (4) Belt Supergroup,North America (Hall and Veizer, 1996); (5) Bakal Formation, Southern Urals (this work); (6) Kotuikan Formation, Anabar Uplift(Knoll et al., 1995); (7) Lower Yusmastakh Subformation, Anabar Uplift (Knoll et al., 1995); (8) Upper Yusmastakh subformation(Knoll et al., 1995); (9) Society Cliffs Formation, Canada (Kah et al., 2001); (10) Malgina Formation of the Uchur–Maya region,Linok and Sukhaya Tunguska formations of the Turukhansk Uplift in Siberia (Bartley et al., 2001); (11) Lakhanda Group of theUchur–Maya region and Burovaya, Shorikha, Miroedikha, and Turukhansk formations of the Turukhansk region in Siberia (Bartleyet al., 2001).



India, Australia, the Urals, and Rocky Mountains ofNorth America, which are considered above, imply thatδ13ë variations in seawater approached ca. 1600 Maago the level typical of the Middle Riphean. Accord-ingly, the δ13ë global stasis, when values of this param-eter were within the range of 0 ± 1‰, terminated beforethe commencement of the Early Riphean.

The 87Sr/86Sr ratios in seawater evaluated for theBurzyan paleobasin corresponded to 0.70460–0.704801550 ± 30 Ma ago and to 0.70456–0.70481 1430 ±30 Ma ago. Our results are consistent with the Sr-isoto-pic characteristics of the Early Riphean seawater estab-lished in the other regions, where the least altered car-bonates of the Lower Riphean have been studied. Theage of the studied rocks is inferred from isotopic-geo-chronological data of variable validity and ranges in gen-eral from 1630 to 1380 Ma. The established 87Sr/86Srratios correspond to 0.70465 in the dolomites of theKyutingda Formation in the Olenek Uplift (Gorokhovet al., 1995), to 0.70478–0.70501 in the dolomites of theKotuikan Formation in the Anabar Uplift (Pokrovskiiand Vinogradov, 1991), to 0.70484–0.70514 in the lime-stones of the Newland Formation of the Belt Supergroup(Hall and Veizer, 1996), and to 0.70460–0.70494 in thelimestones of the Semri Group of central India (Rayet al., 2003).

Hence, the available Sr-isotopic data evidence thepredominant influx of mantle derived components intosea basins of the Early Riphean (Kuznetsov et al., 2008,and references therein). The low and slightly variablevalues of the 87Sr/86Sr initial ratio, which are character-istic of the Lower Riphean carbonates, form the ascend-ing trend at ca. 1250–1220 Ma in connection with theElzevirian orogeny (Semikhatov et al., 2002). The low87Sr/86Sr ratio in the Early Riphean Ocean controlled theisotopic specifics of marine chemogenic sediments atthat time. However, the transition from a slightly fluc-tuating and low 87Sr/86Sr ratio of the Early Riphean tothe higher and more variable values of the late MiddleRiphean is too gradual, and gradients of the valuesunder consideration can hardly be used for substantia-tion of the boundary between the Lower and MiddleRiphean. Nevertheless, the 87Sr/86Sr initial ratio in theLower Riphean carbonates (0.7045–0.7050) percepti-bly differs from the ratio ranging approximately from0.7052 to 0.7072 and characterizing rocks of the termi-nal Middle and Upper Riphean (Gorokhov et al., 1995;Kuznetsov et al., 1997, 2003·, 2006; Bartley et al.,2001; Semikhatov et al., 2002).

In conclusion, we would like to note the specialimportance of the Lower Riphean type succession ofthe Southern Urals in the Late Proterozoic geologicalrecord. Multidisciplinary investigation of carbonaterocks from the Burzyan Group of the Urals and itsequivalents in Russia provided valuable informationabout the biotic and isotopic-geochemical specifics ofthe Early Riphean sea basins. As a result, we revised theage constraints postulated for the global stasis of the

carbonate carbon isotopic composition and obtainednew data confirming the significance of the mantle Srinflux into the oceans of the Early Riphean time.

ACKNOWLEDGMENTS

We are grateful to A.K. Khudolei and A.M. Larin fortheir constructive comments on our work. We alsothank I.V. Kislova, who analyzed trace element concen-trations in carbonates, V.A. Polyakov for the determina-tion of carbon and oxygen isotope composition in thestudied samples, G.V. Konstantinova and E.P. Kutyavinfor analysis of Sr concentrations in carbonate fractions.We also thank M.T. Krupenin for his assistance in orga-nizing our fieldwork and collecting samples.

The work, supported by the Russian Foundation forBasic Research (project nos. 05-05-65290, 05-05-65329,07−05-01107, and 08-05-00429), is carried out as partof the Presidium of the Russian Academy of Sciences’Priority Program no. 15 and of the Program for BasicResearch no. 8 of the Earthscience Division of the Rus-sian Academy of Sciences.

Reviewers A.K. Khudoleiand M.A. Fedonkin

REFERENCES

1. L. V. Anfimov, Lithogenesis in the Riphean SedimentarySuccessions of the Bashkirian Meganticlinorium(S. Urals), (Uralian Division RAS, Yekaterinburg,1997) [in Russian].

2. L. V. Anfimov, B. D. Busygin, and L. E. Demina, TheSatka Magnesite Deposits (Nauka, Moscow, 1983)[in Russian].

3. L. V. Anfimov, A. M. Sul’man, and V. G. Petrishcheva,“Strontium in the Riphean Carbonate Rocks,” inGeochemistry of Sedimentary Formations of the Urals(UNTs AN SSSR, Sverdlovsk, 1987), pp. ë. 22–28[in Russian].

4. J. L. Banner and G. N. Hanson, “Calculation of Simul-taneous and Trace Element Variations during Water–Rock Interaction with Applications of CarbonateDiagenesis,” Geochim. Cosmochim. Acta 54 (11),3123–3137 (1990).

5. J. K. Bartley, L. C. Kah, J. L. McWilliams, A. F. Stag-ner, “Carbon Isotope Chemostratigraphy of the MiddleRiphean Type Section (Avzyan Formation, SouthernUrals, Russia): Signal Recovery in Fold-and-ThrustBelt,” Chem. Geol. 237 (1–2), 211–232 (2007).

6. J. K. Bartley, M. A. Semikhatov, A. J. Kaufman et al.,“Global Events across the Mesoproterozoic–Neoprot-erozoic boundary: C and Sr Isotopic Evidence fromSiberia,” Precambrian Res. 111 (1–4), 165–202 (2001).

7. A. M. Belyaev, L. A. Neymark, A. D. Shebanov, and A.M. Larin, “Age and Origin of Mafic Xenoliths fromRapakivi Granites of the Berdiaush massif (S. Urals,Russia),” in Proceedings of the 7th Intern. Symposiumon Rapakivi Granites and Related Rocks, Helsinki1996,p. 6.

598


SEMIKHATOV et al.

8. U. Brand and J. Veizer, “Chemical Diagenesis of a Mul-ticomponent Carbonate System – 2. Stable Isotopes,”J. Sediment. Petrol. 51 (3), 987–997 (1980).

9. M. D. Brasier and J. F. Lindsay, “A Billon Years of Envi-ronmental Stability and the Emergence of Eukaryotes:New Data from Northern Australia,” Geology 26 (6),555–558 (1998).

10. S. Chaudhuri and N. Clauer, “Strontium Isotopic Com-positions and Potassium and Rubidium Contents of For-mation Waters in Sedimentary Basins: Clues to the Ori-gin of the Solutes,” Geochim. Cosmochim. Acta 57 (3),429–437 (1993).

11. J.W. Cowie and M.R.W. Johnson, “Late Precambrianand Cambrian Geological Time Scale. The Chronologyof the Geological Record,” J. Geol. Soc. London 10,47–64 (1985).

12. J. I. Drever, The geochemistry of natural water (Pren-tice-Hall, Cambridge, 1982).

13. R. Elmis, M. T. Krupenin, V. I. Bogatov, andN. V. Chaplygina, “The Early–Middle Riphean Age ofMain Generation of Diabase Dykes in the Lower Riph-ean Rocks of the Bakal Region (the Southern Urals),”Proceedings of the 2nd All-Russia Petrographic Con-ference: Petrography across the XXI Century Boundary(Results and Prospects), Vol. 4 (IG Komi NTS URORAN, Syktyvkar, 2000), pp. 228–230 [in Russian].

14. R. E. Ernst, K. L. Buchan, M. A. Hamilton, et al., “Inte-grated Paleomagnetism and U–Pb Geochronology ofMafic Dikes of the Eastern Anabar Shield Region, Sibe-ria: Implications for Mesoproterozoic Paleolatitude ofSiberia and Comparison with Laurentia,” J. Geol. 108(3), 381–401 (2000).

15. R. E. Ernst, V. Pease, V. N. Puchkov et al., “Geochemi-cal Characterization of Precambrian Magmatic Suitesof the Southeastern Margin of the East European Cra-ton, Southern Urals, Russia,” Geol. Collection(IG UNTS RAN, Ufa), No. 5, 119–161 (2006).

16. P. Fritz and D. G. Smith, “The Isotopic Composition ofSecondary Dolomites,” Geochim. Cosmochim. Acta 34(11), 1161–1173 (1970).

17. G. Gaal and R. Gorbatschev, “An outline of the Precam-brian evolution of the Baltic Shield,” Precambrian Res.35 (1), 15–52 (1987).

18. M. I. Garan’, “The Upper Precambrian (Riphean):Stratigraphy,” in Stratigraphy of the USSR, Vol. 12,Pt.1: Geological Description (Nedra, Moscow, 1969),pp. 70–106 [in Russian].

19. M. A. Garris, “Geochronological Scale of the Urals andMain Stages of Its Evolution in the Precambrian andPaleozoic (Implication of the K–Ar Dating),” in Abso-lute Age of Geological Formations (Nauka, Moscow,1964), pp. 128–156 [in Russian].

20. I. M. Gorokhov, M. A. Semikhatov, E. R. Drubetskoi,et al., “The Rb–Sr and K–Ar Age of Sedimentary Geo-chronometers from the Lower Riphean of the AnabarMassif,” Izv. Akad. Nauk SSSR, Ser. Geol., No. 7, 17–32 (1991).

21. I. M. Gorokhov, M. A. Semikhatov, A. V. Baskakov, etal., “Sr Isotopic Composition in Riphean, Vendian, andLower Cambrian Carbonates from Siberia,” Stratigr.Geol. Korrelyatsiya 3 (1), 3–33 (1995) [Stratigr. GeolCorrelation 3 (1), 1–29 (1995)].

22. F. M. Gradstein, I. G. Ogg, A. G. Smith, et al., “A NewGeologic Time Scale, with Special Reference to Pre-cambrian and Neogene,” Episodes 27 (2), 83–100(2004a).

23. Gradstein F.M., Ogg I.G., Smith A.G. et al., A GeologicTime Scale, 2004 (Cambridge Univ. Press, Cambridge,2004b).

24. S. Hall and J. Veizer, “Geochemistry of PrecambrianCarbonates: VII. Belt Supergroup, Montana and Idaho,USA,” Geochim. Cosmochim. Acta 60 (4), 667–677(1996).

25. W. B. Harland, R. L. Armstrong, A. V. Cox, et al.,A Geologic Time Scale, 1989 (Cambridge Univ. Press,Cambridge, 1990).

26. H. J. Hofmann and J.W. Schopf, “Early ProterozoicMicrofossils,” in Earth’s Earliest Biosphere: Its Originand Evolution (Princeton Univ. Press, Princeton, 1983),pp. 321–360.

27. H. Hostenberger and M. Herman, “KH-2, the New13C/12C and 18O/16O Standard for the Mass Spectrome-try of Geological Samples,” Geokhimiya, No. 12,1939–1940 (1984).

28. L. C. Kah, D. C. Crawford, J. K. Bartley, et al., C- andSr-isotope Chemostratigraphy as a Tool for VerifyingAge of Riphean Deposits in the Kama–Belaya Aulaco-gen, the East European Platform,” Stratigr. Geol. Korre-lyatsiya 15 (1), 15–34 (2007) [Stratigr. Geol. Correla-tion 15 (1), 12–29 (2007)].

29. L. C. Kah, T. W. Lyons, and J. T. Chesley, “Geochemis-try of a 1.2 Ga Carbonate–Evaporite Succession, North-ern Baffin and Bylot Islands: Implications for Mesopro-terozoic Marine Evolution,” Precambrian Res. 111(1/4), 203–234 (2001).

30. A. J. Kaufman and A. H. Knoll, “Neoproterozoic Vari-ations in the Carbon Isotopic Composition of Seawater:Stratigraphic and Geochemical Implications,” Precam-brian Res. 73 (1/4), 27–49 (1995).

31. B. M. Keller and K. O. Krats, “General StratigraphicScale of the Precambrian in the USSR and Approachesto Its Perfection,” in Main Problems of Subdividing thePrecambrian in the USSR (Nauka, Leningrad, 1979),pp. 127–146 [in Russian].

32. B. M. Keller, K. O. Krats, F. P. Mitrofanov, et al.,“Progress in Elaborating the General StratigraphicScale of the Precambrian in the USSR,” Izv. Akad.Nauk SSSR, Ser. Geol., No. 11, 16–21 (1977).

33. A. H. Knoll and M. A. Semikhatov, “The Genesis andTime Distribution of Two Distinctive Proterozoic Stro-matolite Microstructure,” Palaios 13 (4), 408–422(1998).

34. A. H. Knoll, A. J. Kaufman, and M. A. Semikhatov,“The Carbon-Isotope Composition of Proterozoic Car-bonates: Riphean Successions from Northwestern Sibe-ria (Anabar Massif, Turukhansk Uplift),” Amer. J. Sci.295 (6), 823–850 (1995).

35. V. I. Kozlov, A. A… Krasnobaev, I. N. Larionov, et al.,Lower Riphean of the Southern Urals (Nauka, Moscow,1989) [in Russian].

36. V. I. Kozlov, A. F. Veis, N. D. Sergeeva, et al., “Lithology,Petrography, and Microfossils of the Upper CambrianDeposits Recovered by the Parametric Borehole 133Azino-Pal’nikovskaya in Udmurtia,” in Ezhegodni-



1996 (Inst. Geol. UNTs RAS, Ufa, 1998), pp. 63–70[in Russian].

37. V. A. Komar, Stromatolites from the Upper Precam-brian Deposits in North Siberian Platform and TheirStratigraphic Significance (Nauka, Moscow, 1966)[in Russian].

38. V. A. Komar, M. E. Raaben, and M. A. Semikhatov, TheRiphean Conophyton Forms in the USSR and TheirStratigraphic Significance (Nauka, Moscow, 1965)[in Russian].

39. A. A. Krasnobaev, Zircon as Indicator of GeologicalProcesses (Nauka, Moscow, 1986) [in Russian].

40. A. A. Krasnobaev, G. B. Fershtater, A. I. Stepanov,et al., “Petrology and Rb–Sr Geochronology of the Ber-dyaush Rapakivi Massif,” Izv. Akad. Nauk SSSR,Ser. Geol., No. 1, 21–38, (1981).

41. A. A. Krasnobaev, E. V. Bibikova, A. I. Stepanov, et al.,“Geochronology and Genesis of the Berdyaush Massif(the Urals),” Izv. Akad. Nauk SSSR, Ser. Geol., No. 3,3–23 (1984)

42. A. A. Krasnobaev, E. V. Bibikova, Yu. L. Ronkin, et al.,“Geochronology of Volcanics from the Ai Formationand Isotopic Age of the Riphean Lower Boundary,”Izv. Akad. Nauk SSSR, Ser. Geol., No. 6, 25–40,(1992).

43. M. T. Krupenin, Formation Environment of Siderites inthe Lower Riphean Bakal Formation (the SouthernUrals) (Uralian Division RAS, Yekaterinburg, 1999)[in Russian].

44. M. T. Krupenin and V. Prochaska, “The EvaporiteNature of Fluid Inclusions in Sparry Magnesites of theSatka Type,” Dokl. Ross. Akad. Nauk 403 (5) 661–663(2005) [Dokl. Earth Sciences 403A (6), 838–840(2005)].

45. M. T. Krupenin, A. V. Maslov, M. V. Rykus, and V. I.Sukachev, “New Data on Corg. Concentration in Shalesof the Lower and Middle Riphean from the SouthernUrals,” in Ezhegodnik-1992 (IGG URO RAS, Yekater-inburg, 1993), pp. 19–20 [in Russian].

46. I. N. Krylov, Columnar Branching Stromatolites inRiphean Deposits of the Southern Urals and Their Sig-nificance for the Upper Precambrian Stratigraphy(Nauka, Moscow, 1963) [in Russian].

47. I. N. Krylov, Riphean and Phanerozoic Stromatolites inthe USSR (Nauka, Moscow, 1975) [in Russian].

48. I. N. Krylov and V. N. Sergeev, “Riphean Microfossilsfrom the Kusa Town Area, the Southern Urals,” inStratigraphy, Lithology and Geochemistry of the UpperPrecambrian in the Southern Urals and Cis-Urals(BF AN SSSR, Ufa, 1986), pp. 51–56 [in Russian].

49. A. B. Kuznetsov, I. M. Gorokhov, M. A. Semikhatov etal., “Strontium Isotopic Composition in the Limestonesof the Inzer Formation, Upper Riphean Type Section,Southern Urals,” Dokl. Ross. Akad. Nauk 353 (2) 249–254 (2005) [Dokl. Earth Sciences 353 (2), 319–324(2005)].

50. A. B. Kuznetsov, G. V. Ovchinnikova, I. M. Gorokhovet al., “Sr-Isotope Signature and Pb–Pb Age of theBakal Formation Limestones in the Lower RipheanType Section, the Southern Urals,” Dokl. Ross. Akad.Nauk 391 (6) 794–798 (2003a) [Dokl. Earth Sciences391A (6), 819–822 (2003a)].

51. A. B. Kuznetsov, M. A. Semikhatov, I. M. Gorokhovet al., “Sr Isotope Composition in Carbonates of theKaratau Group, Southern Urals, and Standard Curve of87Sr/86Sr Variations in the Late Riphean Ocean,”Stratigr. Geol. Korrelyatsiya 11 (5), 3–39 (2003b)[Stratigr. Geol. Correlation 11 (5), 415–449 (2003b)].

52. A. B. Kuznetsov, M. T. Krupenin, G. V. Ovchinnikova,et al., “Diagenesis of Carbonate and Siderite Depositsof the Lower Riphean Bakal Formation, the SouthernUrals: Sr Isotopic Characteristics and Pb–Pb Age,”Litol. Polezn. Iskop., No. 3, 227–249 (2005) [Lithol.Miner. Resources 40 (3), 195–215 (2005)].

53. A. B. Kuznetsov, M. A. Semikhatov, A. V. Maslov, et al.,“New Data on Sr- and C-Isotopic Chemostratigraphy ofthe Upper Riphean Type Section (Southern Urals),”Stratigr. Geol. Korrelyatsiya 14 (6), 25–53 (2006)[Stratigr. Geol. Correlation 14 (6), 602–628 (2006)].

54. A. B. Kuznetsov, G. V. Ovchinnikova, M. A. Semi-khatov et al., “The Sr Isotopic Characterization and Pb–Pb Age of Carbonate Rocks from the Satka Formation,the Lower Riphean Burzyan Group of the SouthernUrals,” Stratigr. Geol. Korrelyatsiya 16 (2), 16–34(2008) [Stratigr. Geol. Correlation 16 (2), 120–137(2008)].

55. A. M. Larin, Yu. V. Amelin, L. A. Neymark, and R. Sh.Krymsky, “The Origin of the 1.73–1.70 Ga AnorogenicUlkan Volcano-Plutonic Complex, Siberian Platform,Russia: Inferences from Geochronological, Geochemi-cal, and Nd–Sr–Pb Isotopic Data,” An. Acad. Brazil. 69(3), 295–312 (1997).

56. J. F. Lindsay and M. D. Brasier, “A Carbon Isotope Ref-erence Curve for ca. 1700–1575 Ma, McArthur andMount Isa Basins, Northern Australia,” PrecambrianRes. 99 (3/4), 271–308 (2000).

57. J. F. Lindsay and M. D. Brasier, “Did Global TectonicDrive Early Biosphere Evolution? Carbon IsotopeRecord from 2.6 to 1.9 Ga Carbonate of Western Aus-tralia Basins,” Precambrian Res. 114 (1–2), 1–34(2002).

58. A. V. Maslov, “Paleogeography of the Burzyanian inStratotype Locality,” Litol. Polezn. Iskop., No. 2, 133–149 (1997) [Lithol. Miner. Resources 32 (2), 113–127(1997)].

59. A. V. Maslov, M. T. Krupenin, E. Z. Gareev, and L. V.Anfimov, The Riphean in Western Flank of the SouthernUrals (Classical Sections, Sedimentology, Lithogenesis,Minerageny, Natural Geological Monuments), Vol. 1(IGG URO RAS, Yekaterinburg, 2001) [in Russian].

60. A. V. Maslov, Yu. L. Ronkin, M. T. Krupenin, et al.,“The Lower Riphean Fine-Grained AluminosilicateClastic Rocks of the Bashkir Anticlinorium in theSouthern Urals: Composition and Evolution of TheirProvenance,” Geokhimiya, No. 6, 649–669 (2004)[Geochemistry International 42 (6), 561–578 (2004)].

61. J. A. McKenzie, “Holocene Dolomitization of CalciumCarbonate Sediments from the Coastal Sabkhas of AbuDabi, U.A.F.: A Stable Isotope Study,” J. Geol. 89 (2),185–198 (1981).

62. G. V. Ovchinnikova, I. M. Vasil’eva, M. A. Semikhatovet al., “U-Pb Systematics on Proterozoic CarbonateRocks: The Inzer Formation of the Upper Riphean Stra-totype (Southern Urals) Stratigr. Geol. Korrelyatsiya 6

600


SEMIKHATOV et al.

(4), 20–31 (1998) [Stratigr. Geol. Correlation 6 (4),336–347 (1998)].

63. G. V. Ovchinnikova, I. M. Vasil’eva M. A. Semikhatovet al., “The Pb–Pb Trail Dating of Carbonates withOpen U–Pb Systems: The Min’yar Formation of theUpper Riphean Stratotype, Southern Urals,” Geol. Kor-relyatsiya 8 (6), 3–19 (2000) [Stratigr. Geol. Correla-tion 8 (6), 529–543 (2000)].

64. R. W. Page, T. J. Griffin, I. M. Tyler, and S. Sheppard,“Geochronological Constraints on Tectonic Models forAustralian Palaeoproterozoic High-K Granites,”J. Geol. Soc. London. 158 (3), 535–545 (2001).

65. V. I. Petrov, A. E Shalaginov, and B. N. Punegov, “Phos-phorite in the Lower Riphean Satka Formation of theSouthern Urals,” in Proceedings of the Conference onGeneral Problems of the Riphean Stratigraphy andGeological History in North Eurasia (IGG URO RAS,Yekaterinburg, 1995), pp. 78–79 [in Russian].

66. P. Yu. Petrov and M. A. Semikhatov, “ Specific Devel-opment of the Late Riphean Siliciclastic–CarbonateShelf: the Derevnya Formation of the TurukhanskUplift, Siberia, and Lithostratigraphic Equivalents,”Stratigr. Geol. Korrelyatsiya 13 (3), 26–52 (2005)[Stratigr. Geol. Correlation 13 (3), 242–265 (2005)].

67. K. A. Plumb, “New Precambrian Time Scale,” Episodes14 (2), 139–140 (1991).

68. K. A. Plumb and H. L. James Subdivision of Precam-brian Time: Recommendation and Suggestion by theSubcommission on Precambrian Stratigraphy,” Precam-brian Res. 32 (1), 65–92 (1986).

69. B. G. Pokrovskii and V. I, Vinogradov, “Strontium,Oxygen and Carbon Isotope Composition in the UpperPrecambrian Carbonates from Western Flank of theAnabar Uplift (the Kotuikan River),” Dokl. Akad. NaukSSSR 320 (5), 1245–1250 (1991).

70. B. Rasmussen, P. K. Bose, S. Sarkar, et al., “1.6 Ga U–Pb Zircon Ages for the Chorhat Sandstone, LowerVindhyan, India: Possible Implications for Early Evolu-tion of Animals,” Geology 30, 103–106 (2002).

71. J.S. Ray, J. Veizer, and W. J. Davis, “C, O, Sr and Pb Iso-tope Systematics of Carbonate Sequences of theVindhyan Supergroup, India: Age, Diagenesis, Correla-tions and Implications for Global Events,” PrecambrianRes. 121 (1/2), 103–140 (2003).

72. Resolution of the All-Union Stratigraphic Conferenceon General Problems of the Precambrian Subdivisionin the USSR (Nauka, Leningrad, 1979) [in Russian].

73. “Resolution of the Third All-Russian Meeting on Gen-eral Problems of Precambrian Stratigraphy” Stratigr.Geol. Korrelyatsiya 9 (3), 101–106 (2001) [Stratigr.Geol. Correlation 9 (3), 304–308 (2001)].

74. Yu. L. Ronkin and O. P. Lepikhina, “Petrogenesis andSm–Nd Isotopic Age of Rapakivi Granite from the Ber-dyaush Massif (the Southern Urals),” in Proceedings ifthe Intern. Sci. Conference on Ophiolites: Geology,Petrology, Metallogeny and Geodynamics (IGG URORAN, Yekaterinburg, 2006), pp. 322–325 [in Russian].

75. Yu. L. Ronkin and O. P. Lepikhina, “New U–PbSHRIMP-II Data on Age of Volcanics from the Ai For-mation, the Burzyan Group of the Bashkirian Meganti-clinorium, and Age Problem of the Riphean Base,” inProceedings of the 5th All-Russian Lithologic Confer-

ence: Sedimento- and Lithogenesis and Their Evolutionin the Earth’s History (IGG URO RAN, Yekaterinburg,2008), pp. 203–207 [in Russian].

76. Yu. L. Ronkin, D. N. Matukov, S. L. Presnyakov, et al.,“The U–Pb SHRIMP In-Situ Dating of Zircons fromNepheline Syenite of the Berdyaush Massif (the South-ern Urals),” Litosfera, No. 1, 135–142 (2005).

77. Yu. L. Ronkin, A. V. Maslov, A. P. Kazak, et al., “TheLower–Middle Riphean Boundary in the SouthernUrals: New Isotopic U–Pb (SHRIMP II) ConstraintsDokl. Akad. Nauk 415 (3), 370–376 (2007) [DokladyEarth Sci. 415A (6), 835–841 (2007)].

78. G. M. Ross and M. Villeneuve, “Provenance of theMesoproterozoic (1.45 Ga) Belt Basin (Western NorthAmerica): Another Piece in the Pre-Rodinia Paleogeo-graphic Puzzle,» Bull. Geol. Soc. Am. 115 (10), 1191–1217 (2003).

79. M. V. Rykus, N. I. Andreev, V. P. Murkin, et al., The Pre-cambrian Carbonaceous Deposits in the SouthernUrals (IG UNTs RAN, Ufa, 1993) [in Russian].

80. L. I. Salop, General Stratigraphic Scale of the Precam-brian (Nedra, Leningrad, 1973) [in Russian].

81. L. I. Salop and G. A. Murina, “Age of the BerdyaushRapakivi Pluton and Problem of GeochronologicalBoundaries of the Lower Riphean,” Sov. Geol., No. 6,15–27 (1970).

82. J. W. Sears, K. R. Chamberlain, and S. N. Buckly,“Structural and U–Pb Geochronological Evidence for1.47 Ga Rifting in the Belt Basin, Western Montana,”Can. J. Earth Sci. 35, 467–475 (1998).

83. M. A. Semikhatov, “Some Carbonate Stromatolites ofthe Aphebian from the Canadian Shield,” in LowerBoundary of the Riphean and Stromatolites of the Aph-ebian (Nauka, Moscow, 1978), pp. 111–147 [in Rus-sian].

84. M. A. Semikhatov, “New Stratigraphic Scale of theUpper Precambrian in the USSR: Analysis and the Les-sons,” Izv. Akad. Nauk SSSR, Ser. Geol., No. 11, 5–22(1979).

85. M. A. Semikhatov, General Problems of ProterozoicStratigraphy in the USSR, Soviet Sci. Rev. Sec. G. Geol.Rev. (Harwood Acad. Publ., New York, 1991).

86. M. A. Semikhatov, “Current Concepts of General Sub-division of the Precambrian: An Analysis,” Izv. Akad.Nauk SSSR, Ser. Geol., No. 8, 3–13 (1991).

87. M. A. Semikhatov, “General Subdivision of the Pre-cambrian: Comparison of the Latest Scales,” Stratigr.Geol. Korrelyatsiya 1 (1), 6–20 (1993) [Stratigr. Geol.Correlation 1 (1), 1–15 (1993)].

88. M. A. Semikhatov and M. E. Raaben, “ Dynamics of theGlobal Diversity of Proterozoic Stromatolites. Article 1:North Eurasia, China, and India Stratigr. Geol. Korre-lyatsiya 2 (6), 10–32 (1994) [Stratigr. Geol. Correlation2 (6), 12–29 (1994)].

89. M. A. Semikhatov and M. E. Raaben, “Dynamics of theGlobal Diversity of Proterozoic Stromatolites. Article 2:Africa, Australia, North America, and General Synthe-sis Stratigr. Geol. Korrelyatsiya 4 (1), 26–54 (1996)[Stratigr. Geol. Correlation 4 (1), 24–50 (1996)].

90. M. A. Semikhatov and S. N. Serebryakov, SiberianHypostratotype of the Riphean (Nauka, Moscow, 1983)[in Russian].



91. M. A. Semikhatov, A. B. Kuznetsov, I. M. Gorokhov etal., “Low 87Sr/86Sr Ratios in Seawater of the Grenvilleand post-Grenville Time: Determining Factors,”Stratigr. Geol. Korrelyatsiya 10 (1), 1–41 (2002)[Stratigr. Geol. Correlation 10 (1), 1–41 (2002)].

92. M. A. Semikhatov, A. B. Kuznetsov, V. N. Podkovyrov,et al., “The Yudoma Group of Stratotype Area: C-iso-tope Chemostratigraphic Correlations and Yudomian–Vendian Relation,” Stratigr. Geol. Korrelyatsiya 12 (5),3–29 (2004) [Stratigr. Geol. Correlation 12 (5), 435–459 (2004)].

93. M. A. Semikhatov, K. A. Shurkin, E. M. Aksenov, et al.,“New Stratigraphic Scale of the Precambrian in theUSSR,” Izv. Akad. Nauk SSSR, Ser. Geol., No. 4, 3–13(1991).

94. V. N. Sergeev, Silicified Microfossils of the Precam-brian: Nature, Classification, and Biostratigraphic Sig-nificance (GEOS, Moscow, 2006) [in Russian].

95. V. N. Sergeev and Lee Seong-Joo, “Microfossils fromCherts of the Middle Riphean Svetlyi Formation, theUchur–Maya Region of Siberia, and Their StratigraphicSignificance,” Stratigr. Geol. Korrelyatsiya 9 (1), 3–13(2001) [Stratigr. Geol. Correlation 9 (1), 1–10 (2001)].

96. V. N. Sergeev and Lee Seong-Joo, “New Data on Silic-ified Microfossils from the Satka Formation of theLower Riphean Stratotype, the Urals,” Stratigr. Geol.Korrelyatsiya 12 (1), 3–26 (2004) [Stratigr. Geol. Cor-relation 12 (1), 1–21 (2004)].

97. S. N. Serebryakov, Peculiarities in Formation and Dis-tribution of Riphean Stromatolites in Siberia (Nauka,Moscow, 1975) [in Russian].

98. M. S. Shvetsov, Petrography of Sedimentary Rocks(Gosgeoltekhizdat, Moscow, 1958) [in Russian].

99. S. Sindern, R. Hetzel, B. A. Schulte, et al., “ProterozoicMagmatic and Tectonometamorphic Evolution of theTaratash Complex, Central Urals, Russia,” Int. J. EarthSci. 94, 319–335 (2005).

100. S. Sindern, Yu. L. Ronkin, R. Hetzel, et al., “TheTaratash and Aleksandrovsk Metamorphic Complexes:The P–T Constraints,” in Ezhegodnik-2005 (IGG URORAN, Yekaterinburg, 2006), pp. 322–330 [in Russian].

101. Stratigraphic Code, Second Edition Supplemented(MSK, St. Petersburg, 1992) [in Russian].

102. Stratigraphic Scheme of Riphean and Vendian Depositsin the Volga–Ural Region, Explanatory Notes(IG UNTs RAS, Ufa, 2000) [in Russian].

103. Stratotype of the Riphean: Paleontology, Paleomag-netism (Nauka, Moscow, 1982) [in Russian].

104. Stratotype of the Riphean: Stratigraphy, Geochronology(Nauka, Moscow, 1983) [in Russian].

105. A. F. Veis and N. G. Vorob’eva, “The Riphean and Ven-dian Microfossils of the Anabar Massif,” Izv. Akad.Nauk SSSR, Ser. Geol., No. 1, 114–130 (1992).

106. A. F. Veis and N. G. Vorob’eva, “Microfossils from theRiphean Hypostratotype of Siberia (the Omakhta,Kandyk, and Ust-Kirba Microbiotas),” Stratigr. Geol.Korrelyatsiya 10 (2), 27–54 (2002) [Stratigr. Geol. Cor-relation 10 (2), 125–150 (2002)].

107. A. F. Veis and M. A. Semikhatov, “The Lower RipheanOmakhta Assemblage of Microfossils from East Sibe-ria: Composition and Formation,” Izv. Akad. NaukSSSR, Ser. Geol., No. 5, 36–55 (1989).

108. A. F. Veis, E. V. Kozlova, and N. G. Vorob’eva,“Organic-Walled Microfossils from the Riphean TypeSection,” Izv. Akad. Nauk SSSR, Ser. Geol., No. 9, 20–36 (1990).

109. A. F. Veis, E. V. Kozlova, N. G. Vorob’eva, and LeeSeong-Joo, “Geochronological and BiostratigraphicApproaches to Reconstruction of Precambrian BiotaEvolution: New Finds of Microfossils in Riphean Sec-tions on the Western Slope of the Anabar Uplift,” Dokl.Ross. Akad. Nauk 378 (4), 511–517 (2001) [DokladyEarth Sciences 378 (4), 413–419 (2001)].

110. J. Veizer and J. Hoefs, “The Nature of 18O/16O and13C/12C Secular Trends in Sedimentary CarbonateRocks,” Geochim. Cosmochim. Acta. 40 (11), 1387–1395 (1976).

111. J. Veizer, K. A. Plumb, R. N. Clayton, et al., “Geochem-istry of Precambrian carbonates: Late Paleoproterozoicseawater,” Geochim. Cosmochim. Acta. 56 (6), 2487–2501 (1992).

112. F. Ya. Vlasov, “Precambrian Stromatolites from theSatka Formation of the Southern Urals,” in Paleonto-logical Data on the Middle Paleozoic of Siberia and theUrals (UNTs AN SSSR, Sverdlovsk, 1977), pp. 101–128 [in Russian].

113. M. R. Walter, I. N. Krylov, and M. D. Mutz, “Stromato-lites from Middle and Late Proterozoic Sequences in theMcArthur and Georgina Basins and the Mount Isa Prov-ince, Australia,” Alcheringa 12, 79–106 (1988).

114. T. V. Yankauskas, N. S. Mikhailova, T. N. German,et al., “Precambrian Microfossils in the USSR (Nauka.Leningrad, 1989) [in Russian].

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