δ 13C and δ 18O values of Triassic brachiopods and carbonate rocks as proxies for coeval seawater...

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=Palaeogeography, Palaeoclimatology, Pa
y13C and y18O values of Triassic brachiopods and carbonate rocks

as proxies for coeval seawater and palaeotemperature

Christoph Korte a,*,1, Heinz W. Kozur b, Jan Veizer a,c

aInstitut fur Geologie, Mineralogie und Geophysik, Ruhr-Universitat, 44801 Bochum, GermanybRezsu u. 83, H-1029 Budapest, Hungary

cOttawa-Carleton Geoscience Center, University of Ottawa, Ottawa, Ontario, Canada K1N 6N5

Received 22 March 2004; received in revised form 28 April 2005; accepted 20 May 2005

Abstract

A dataset of 160 isotope y13C and y18O values from Anisian, Ladinian, Carnian and Rhaetian articulate brachiopod shells,

complemented by 158 carbon and oxygen isotope values from whole rock carbonates, define the first continuous stable isotope

baseline trends for the Triassic seawater. The carbon isotope data suggest the existence of several short-term high amplitude

excursions in the Early Triassic, followed by a predominance of values around 0.5F1x during the Middle Triassic, a rise to

~3.5x during the Carnian, plateau at this level during the Late Carnian to Early Norian, and a 1.5x decline in the Middle

Norian to values around 2x during the Late Norian–Rhaetian interval. The causation scenarios for these rapid oscillations are at

present equivocal, but may in part reflect a biological instability of the carbon cycle following the recovery from the end-

Permian extinction event and/or an input of bmantleQ-derived CO2 from enhanced volcanic activity.

The y18O values from well-preserved brachiopods from the Tethyan realm range from �3.9 to �0.6x V-PDB. These

values require open marine Triassic seawater y18O values close to 0x V-SMOW for the calculated temperatures to be within

the range of tolerance of the coexisting reef-building corals. A 2x y18O increase within the uppermost Cordevolian and

early Julian suggests either a distinct temperature decline in the Southern Alps during that time interval, an increase in

seawater salinity, or their combination. Oxygen isotope values for the Muschelkalk brachiopods range between �6.2 and

�2.0x and likely reflect a strong influx of meteoric waters into the Germanic Basin that shifted the oxygen isotopes to

more negative values.

D 2005 Elsevier B.V. All rights reserved.

Keywords: Carbon isotopes; Oxygen isotopes; Triassic; Brachiopods; Carbonates

0031-0182/$ - see front matter D 2005 Elsevier B.V. All rights reserved.

doi:10.1016/j.palaeo.2005.05.018

* Corresponding author. Fax: +44 1865 272072.

E-mail address: [email protected] (C. Korte).1 Present address: Department of Earth Sciences, University of

Oxford, Parks Road, Oxford, OX1 3PR, UK.

1. Introduction

Carbon and oxygen isotopic composition of sea-

water during geologic history (Baertschi, 1957; Clay-

ton and Degens, 1959; Degens and Epstein, 1962;

laeoecology 226 (2005) 287–306

C. Korte et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 226 (2005) 287–306288

Weber, 1967; Veizer and Hoefs, 1976; Veizer et al.,

1999) could have been influenced by a combination of

factors, such as climate, waxing and waning of con-

tinental glaciers, burial and re-oxidation of organic

matter, bioproductivity, tectonic activity or changing

oceanic circulation patterns (Shackleton and Opdyke,

1973; Shackleton, 1977; Railsback, 1990; Schi-

dlowski and Aharon, 1992; Grossman, 1994; Kump

and Arthur, 1999). In view of the absence of ancient

seawater samples, its evolution can be reconstructed

only through studies of proxies in carrier phases that

have precipitated from seawater. Examples of such

carriers are articulate brachiopod shells for y13C and

y18O (Compston, 1960; Lowenstam, 1961; Popp et

al., 1986; Veizer et al., 1986, 1999), phosphatic con-

odonts for y18O (Wenzel et al., 2000; Joachimski and

Buggisch, 2002; Korte et al., 2004a) and whole rock

carbonates for y13C (Jenkyns et al., 1994; Weissert et

al., 1998). A prerequisite for deciphering the secular

trends in seawater isotopic composition is that the

signal is free, as much as possible, from diagenetic

and metamorphic overprint (Veizer, 1983a,b). Articu-

late brachiopods with shells composed of low-Mg-

calcite (LMC) are particularly suitable for such studies

because this CaCO3 phase is the most resistant to

diagenetic alteration, thus minimizing the resetting

of the primary geochemical signals (Popp et al.,

1986; Wadleigh and Veizer, 1992; Grossman et al.,

1993, 1996). Such bbest preservedQ brachiopod sig-

nals can, theoretically, be utilized for deciphering the

temperature (Epstein et al., 1953) of ancient seawater.

In this study, the degree of preservation of brachio-

pod shells was monitored by trace element analyses,

optical microscopy, scanning electron microscopy

(SEM) as well as by cathodoluminescence (CL). Tak-

ing into account that none of these screening techni-

ques is foolproof, a combination of all approaches has

to be utilized for selection of the bbest preservedQsamples.

Micritic whole rock carbonates and marine

cements provide an additional useful proxy for sea-

water y13C because for carbon the low water–rock

ratio of the diagenetic system fixes the y13C of the

precursor into the successor phase, the diagenetic

calcite (Veizer, 1983b; Weissert, 1989; Marshall,

1992; Grossman, 1994).

In addition to changing seawater composition and

diagenetic overprint, biological fractionation of iso-

topes and elements may also play a role in distorting

the proxy signals. Some marine organisms precipitate

their calcareous shells in equilibrium with the ambient

seawater, while others, such as corals, incorporate

tracers in clear disequilibrium. Such bvital effectsQusually lead to lower y13C and y18O in biogenic

carbonates (Urey et al., 1951; Wefer and Berger,

1991), supposedly due to kinetic fractionation effects

(McConnaughey, 1989a,b). Adkins et al. (2003), on

the other hand, argued that isotopic fractionation is

due to a pH gradient between ambient water and

intracellular fluid, the latter being the solution from

which calcification proceeds. This gradient is partic-

ularly pronounced at high calcification rates. Such a

mechanism may explain why the carbon and oxygen

isotope values in the primary (that is first formed)

shell layers of brachiopods are more depleted in 18O

and 13C than are the slower growing secondary

(meaning later formed) layers (cf. Carpenter and Loh-

mann, 1995). These secondary LMC layers, utilized

exclusively in this study, appear in modern articulate

brachiopods to have been precipitated in approximate

isotopic equilibrium with the ambient seawater (Low-

enstam, 1961; Brand, 1989; Grossman et al., 1991;

Brand et al., 2003). Departures from isotopic equilib-

rium, if they exist, are usually within F1x (Carpen-

ter and Lohmann, 1995). Ancient examples, such as

some Carboniferous specimens, also appear to con-

form to this pattern, with intraspecies differences

usually less than F1x (Mii et al., 2001). Brachiopod

LMC shells thus appear to have been suitable carrier

phases for isotopic proxies of Phanerozoic seawater, a

proposition well documented by a plethora of earlier

studies (Popp et al., 1986; Veizer et al., 1986, 1997;

Grossman et al., 1991, 1993; Bruckschen et al., 1999;

Mii et al., 1999; Korte et al., in press). A summary

compilation for the entire Phanerozoic was published

by Veizer et al. (1999).

The goal of this study is to delineate the baseline

y13C and y18O secular trends for the Triassic inter-

val. Such baselines are essential for research prog-

ress in palaeoceanography, diagenesis, mineral

deposit formation, and palaeotemperatures. The

present large set of the new y13C data traces in

detail carbon isotopic composition of seawater for

almost the entire Triassic interval. This secular curve

is based on both brachiopods and whole rock car-

bonates. The situation is more complex for the

C. Korte et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 226 (2005) 287–306 289

oxygen isotope record, because brachiopods are

available only from specific time intervals. In con-

trast to carbon, the y18O of whole rocks cannot be

utilised for filling the gaps because these samples

are susceptible to resetting during diagenesis. More-

over, the brachiopods and whole rock carbonates

originate from different basins, with variable water

depths, temperatures or diagenetic milieus. As a

result, the y18O secular trend is not as continuous

as would be desirable.

2. Geological setting and stratigraphy

The analysed brachiopod and whole rock carbonate

samples originate from the Northern Alps (Kossen,

Hochalm; Austria), Southern Alps (Pufels, St. Cas-

sian, Uomo; Italy), Balaton Highlands (FelsIors,

Koveskal, Pula Sandorhegy; Hungary), Cserhat

Mountains (CsIvar; Hungary), Mescek Mountains

(Szava; Hungary), Sicanian Palaeogeographic Realm

(Palazzo Adriano, Sosio Valley; Sicily, Italy), Western

P

A

N

G

E

A

Palazzo Adriano(Sicily)

Pufels/Bula/BullaSt. Cassian(Southern Alps)

MuschelkalkSea

Kössen, Hochalm(Northern Alps)

Fig. 1. Palaeogeographic distribution of continents and oceans in the Carni

sampling areas. Continental crust (land and shelf seas) in white, rifts in li

Carpathians (Silicka Brezova; Silica Nappe, Slova-

kia) and Germanic Basin (several sections in Ger-

many, Poland, France) (Figs. 1 and 2). The Tethyan

samples were stratigraphically subdivided using the

conodont zonation of Kozur (1997, 1999, 2003a,b),

except for the St. Cassian sections where we used the

standard ammonoid zones (see Urlichs, 1974, 1994;

Broglio Loriga et al., 1999). For the Muschelkalk

(MK)—deposited in a mid-European epicontinental

shallow sea—the biostratigraphy and lithostratigra-

phy is based on Kozur (1974a,b, 1999), Hagdorn

(1991) and Hagdorn et al. (1993). The stratigraphic

assessments of all the above mentioned sections are

available in Korte (1999), but note that in some cases

the biostratigraphic assessments in Appendix A have

been slightly modified (for Silicka Brezova see also

Channell et al., 2003). Additional samples originate

from the Lower Triassic Pufels (Bula/Bulla) section

(Werfen Formation) of the Italian Southern Alps and

their stratigraphic assessment is published in Perri

(1991) and Farabegoli and Perri (1998) (see also

Korte and Kozur, 2005).

T E T H Y S

equator

Silická Brezová(WesternCarpathians)

Köveskál, FelsoörsCsovár (Hungary)

an to Norian (slightly modified after Stampfli and Borel, 2002), with

ght-grey, oceanic crust in darker grey.

M. koessenensis

M. hernsteini

Neogondolella bifurcata

Paragondolellabulgarica

Nicoraella kockeli

Nicoraella germanica

Sweetospathodus kummeli

Neospathodus dieneri

Hindeodus parvus

Isarcicella isarcica

H. postparvus - H. sosioensis -C. carinata

Neospathodus cristagalli

Chengyuania nepalensis

Neospathodus waageni - S. meeki

N. waageni - Scythogondolella milleri

Triassospathodus hungaricus

Icriospathodus collinsoni

Triassospathodus homeri

Triassospathodus triangularis

Triassospathodus sosioensis

Chiosella gondolelloides

Chiosella timorensis

Neogondolella regalis

Neogondolella constricta

Neogondolella mesotriassica

Paragondolella ? trammeri

Budurovignathus truempyi

Budurovignathus hungaricus

Budurovignathus mungoensis

Budurovignathus n. sp.

Budurovignathus diebeli - P. noah

Gladigondolella tethydis - P. noah

Paragondolella carpathica

Paragondolella noah

Carnepigondolella zoae

Epigondolella quadrata

E. triangularis - N. hallstattensis

Mockina medionorica - M. matthewi

Mockina postera- Mockina zapfei

Mockina bidentata

Misikella hernsteini

Misikellaposthernsteini

Misikella ultima

Mid

dle

Tri

assi

c

Aegean

Pelsonian

Illyrian

Scyt

hian

(E

arly

Tri

assi

c)

Ani

sian

Tuvalian

Julian

Fassanian

Longobardian

Conodont Zone / Subzone

Lad

inia

nC

arni

an

Cordevolian

Parvigondolella andrusovi

Gangetian

Gandarian

Smithian(Lower Olenekian)

Spathian(Upper

Olenekian)

Bithyian

Carnepigondolella pseudodiebeli

Tri

assi

c

Nor

ian

Lacian

Alaunian

Lat

e T

rias

sic

Sevatian

Rha

et.

Substage

Stag

e

Seri

es

Syst

em

Pufels(Southern Alps)

Hochalm,Kössen

(Northern Alps)

Palazzo Adriano

Köveskál(Hungary)



Germany

St.Cassian (S.Alps)

Silická Brezová(Western

Carpathians)

Mus

chel

kalk

several localitiesGermany,Poland,France

Keu

per

Sections

Tethys Germanic BasinGroupSub-

group

M. primitiusE. orchardi - N. navicula

tiepoints

199.6 199.6

Kozur2003a2003b

? 205? 207

237

241.2

238.0

247

252.6

251

252.5

238.8

226

Bun

tsan

dste

in

240.5

Csovár(Hungary)

Felsoörs(Hungary)

249

251.6

225

247.0

Ole

neki

anB

rahm

ania

n(I

ndua

n)

(Early Norian)

Middle(mm)

Lower(mu)

(mo)

UpperM.K.

252.6

Fig. 2. Stratigraphy of the studied sections (conodont zones after Kozur, 2003a,b). The radiometric ages for the tie points originate from Gehrels

et al. (1987), Dunning and Hodych (1990), Lehrmann et al. (2002), Mundil et al. (1996, 2001, 2004) and Palfy et al. (2000, 2003). The

interpolated numerical ages of Kozur (2003a, 2003b) are also shown.


The samples were assigned numeric ages in

accord with the time scales of Kozur (2003a,b).

A detailed discussion of Triassic stratigraphic con-

troversies, radiometric ages and time scales is avail-

able in Korte et al. (2003) and Kozur (2003a,b). In

the following discussion, we prefer to use the

stratigraphic nomenclature rather than the numeric

ages.


3. Methods

3.1. Whole rock carbonates

For whole rock carbonate analyses (only homoge-

neous rock material was used, weathered portions or

veins were rejected), powders were drilled from fresh

surfaces and 150 to 450 Ag were reacted in 10 ml

borosilicate exetainers with phosphoric acid after

flushing with He. Generated CO2 was separated

from water vapor and was analysed for y18O and

y13C on a GasBench II linked to a ThermoFinnigan

DeltaplusXL mass spectrometer at the Institut fur Geo-

logie und Palaontologie, Universitat Innsbruck. The

results were calibrated against V-PDB. For further

details of this method see Spotl and Vennemann

(2003).

3.2. Brachiopods

Analytical work on brachiopod shells was per-

formed at the Institut fur Geologie, Mineralogie und

Geophysik of the Ruhr-Universitat, Bochum. For car-

bon and oxygen isotope analyses aliquots of 3 to 6 mg

shell splinters were reacted offline overnight with

100% phosphoric acid. The generated CO2 was ana-

lysed for y18O and y13C on a Finnigan MAT 251

mass-spectrometer and calibrated against V-PDB.

The NBS 19 values for y18O and y13C were

�2.36F0.03x and 1.92F0.08x, respectively. The

values for NBS 20 were �4.36F0.19 for oxygen and

�1.04F0.07x for carbon.

Despite the fact that LMC articulate brachiopods

are relatively resistant to diagenetic alteration, all

samples were tested for their textural preservation

and chemistry. This screening (optical microscopy,

SEM, ICP-AES, CL) was utilized to select the sam-

ples for isotopic studies or to evaluate the isotope data.

Ca, Sr and Mn contents (Appendix A) were quan-

tified using inductively coupled plasma atomic emis-

sion spectroscopy (ICP-AES) on aliquots remaining

after phosphoric acid treatment of the samples for

carbon and oxygen isotope measurements (Coleman

et al., 1989). The detection limits, depending on the

shell size, were approximately 5 ppm. The precision

was relatively within 10%.

One of the most spectacular tools for evaluation of

the degree of shell presentation is cathodolumines-

cence that was performed on thin sections of the

studied samples. Yet even this is not a foolproof

technique, because parts of modern brachiopod shells

that were not yet affected by diagenesis show bright

orange luminescence (Barbin and Gaspard, 1995). At

the same time, clearly meteorically altered ancient

shells can have only an intrinsic luminescence or no

luminescence at all (Rush and Chafetz, 1990; Qing

and Veizer, 1994). For these reasons, the CL-micros-

copy was used only as one additional information in

evaluating the state of sample preservation.

4. Screening of brachiopods

4.1. Microtextural preservation

The secondary layers of the brachiopod shells were

separated and handpicked under a binocular micro-

scope, in order to inspect them for weathered frag-

ments, attached cements or sediment and crack

fillings. Fragments with such features were rejected.

Shell splinters that passed this first inspection were

further screened by scanning electron microscopy. In

this study, mostly fibrous secondary layers of the

brachiopod shells were utilized. Only samples with

smooth fibrous surfaces were classified as excellently

preserved (Fig. 3a–e). The preservation of fibrous

shapes indicates the absence of dissolution/reprecipi-

tation. Some brachiopods – especially samples from

the Muschelkalk – have punctate shell structures (Fig.

3e). Samples with dissolution or recrystallization fea-

tures (Fig. 3f) were avoided, or isotope values of such

shells were considered as altered.

4.2. Trace elements

For modern brachiopods, 5 to 460 ppm Mn and

200 to 1500 ppm Sr concentrations were quoted by

Morrison and Brand (1986) and Brand (1989) as

typical, but later, for a much larger population,

Brand et al. (2003) cited 1 to 199 ppm for Mn and

450 to 1928 ppm for Sr. Popp et al. (1986) reported

concentrations up to 250 ppm Mn and 300 to 3400

ppm Sr for non-luminescent, texturally well-preserved

Paleozoic brachiopods. In this study, samples with

less than 250 ppm Mn and more than 400 ppm Sr

were classified as well preserved (see also Bruckschen

Fig. 3. SEM images of brachiopod shells. Well-preserved low-Mg-calcite secondary layers of (a) MK/Spiriferina-Bank (locality Muhlhausen-

Wern), (b) MK/Meso 17, (c) Cassian 15, (d) Kossen/D 16, and (e) MK/cycloides-Bank g (locality Seemuhle/Vaihingen/Enz). Some analysed

brachiopods have punctate shells and the punctae (p) may be filled by diagenetic cement (e). However, the proportion of this secondary calcite is

low, thus not distorting of the primary isotope signals. For the diagenetically altered shell (f, sample SB 23), the isotope values are considered to

be reset. Bar scale=20 Am.


et al., 1999; Korte et al., 2003, in press). A lower Sr

content for well-preserved Triassic samples, compared

to modern brachiopods, appears reasonable, since

Steuber and Veizer (2002) documented a secular

Phanerozoic Sr/Ca trend which bottomed in the

Permo-Triassic. More than half of the studied brachio-

pods meet the above trace element requirements (Fig.

4). The exceptions are most samples from the Silicka

Brezova sections, which follow an obvious diagenetic

trend of Brand and Veizer (1980), with progressing

alteration accompanied by Mn enrichment and Sr

depletion.

0.1

1

10

200 400 6000

Mn (ppm)

Sr/C

a *

1000

diagenetictrend

all brachiopods except Silická Brezovábrachiopods from Silická Brezová

Fig. 4. Sr/Ca-Mn cross-plot diagram after Brand and Veizer (1980).

The shaded square (Morrison and Brand, 1986) and the dashed line

one (Brand et al., 2003) are the fields for present-day brachiopods.

The square delineated by solid lines represents well-preserved

samples, as classified in this work.

-2

-1

0

1

2

3

4

5

δ 13 C

‰ (

V-P

DB

)

brachiopodsbrachiopods

wrwr (Korte et

aragonitic fo

wr (Korte an

coal gap very rare coal deposits

strong vol-canic activity

Scythian

Anisian Ladin. Carnian

Middle Triassic

2237240.5247252.6 251

B.

wr TJB-tren

altered brac

Olenek.

Fig. 5. Carbon isotope record of Triassic seawater. The brachiopod data s

with more than 250 ppm Mn and/or less than 400 ppm Sr are classified as

utilized.


5. Results

5.1. Carbon isotopes

The carbon isotope values of all brachiopods and

whole rock carbonates are plotted in Fig. 5 (see also

Appendix A). Whole rock y13C data from Korte et al.

(2004a) and Korte and Kozur (2005) for the lower-

most Triassic are also included. The good textural

preservation (optical microscopy and SEM) of the

brachiopod samples argues for retention of the near-

primary carbon isotope record of ancient seawater.

Note, nevertheless, that carbon isotope values for

well (b250 ppm Mn, N400 ppm Sr) and less well

preserved (N250 ppm Mn, b400 ppm Sr) brachiopods

and of whole rock carbonates all follow similar tem-

poral trends. This is true not only for individual

samples but also for the means of stages with bra-

chiopod y13C differing by less than F0.2x (Korte,

1999).

(Mn < 250 ppm; Sr > 400 ppm) (Mn > 250 ppm; Sr < 400 ppm)

al., 2004a)

ssils

d Kozur, 2005)

coal swamps recovered

beginning of thevolcanism in theCentral AtlanticMagmatic Province

Norian

Late Triassic

Rhaetian

199.626 (Ma)206

d (Pálfy et al., 2001)

hiopods (SEM criteria)

et is subdivided by trace element (Mn, Sr) concentrations. Samples

diagenetically altered. The Triassic time scale of Kozur (2003a,b) is


For the Brahmanian, Olenekian and the earliest

Anisian the whole rock carbon isotope data origi-

nate from three regions: Iran, Sicily and the South-

ern Alps (Figs. 5 and 6). While somewhat different

in absolute values, the trends and their magnitudes

are similar (Fig. 6). Note that the Southern Alps

y13C values at Pufels are about 1.5x lower than

those of the contemporaneous counterparts at Aba-

deh, Iran. The earliest Triassic is characterised by

low y13C values, rising from the well-known nega-

tive carbon isotope anomaly in the latest Permian

(e.g. Baud et al., 1989; Korte et al., 2004a,b). In the

Gangetian, the carbon isotope values increase rapid-

ly, by more than 2.5x, for the Pufels and Abadeh

sections. The Abadeh and the Palazzo Adriano

curves decline by more than 1x in the early Gan-

darian, thus defining the first positive y13C peak in

the late Gangetian. From the middle Gandarian to

the late Smithian only the data from the Southern

Alps are available and these show two distinct

-2

-1

0

1

2

3

4

5wr Pufels (Bula, Bulla), Southern Alps (this study and Kwr Palazzo Adriano, Sicily (this study)

δ 13 C

‰ (

V-P

DB

)

Brahmanian

Scythian (Lower Trias

251252.6

wr Abadeh, Iran (Korte et al., 2004a)

Gangetian Gandar.

251.6

Smithian

end of strong activity ofSiberian Trap volcanism

Fig. 6. Detailed carbon isotope record for the Lower Triassic seawater. N

Alps) are about 2x lower than those of the Abadeh section; the trend, ho

positive peaks, in the basal and middle Smithian.

The 3x decrease in the late Smithian indicates,

tentatively, a negative excursion.

The above Early Triassic trends fit well with the

published whole rock carbon isotope records through-

out the Tethyan realm worldwide.

For Chinese sections, Payne et al. (2004) reported a

similar short positive y13C excursion in the late

Gangetian and values between 0 and 1x for most

of the Gandarian. The second short positive excursion

appears at the base of the Smithian, with values up to

8x. However, its correlation with our section is

somewhat uncertain, because these authors showed

that, from the top of the Neospathodus dieneri–Neo-

spathodus cristagalli fauna (definitive Gandarian) to

the base of the Neospathodus waageni fauna (defi-

nitely Smithian), a longer interval without conodonts

is followed by bPlatyvillosusQ (=Eurygnathodus

Staesche). The real Platyvillosus is a lower Spathian

form that occurs together with Icriospathodus collin-

orte and Kozur, 2005)

sic)

Olenekian Anisian

M. Triassic

247 (Ma)249

Spathian Aegean

beginning ofdistinct faunaland floral recovery

strongfaunal andfloralrecovery

ote that the carbon isotope values of the section at Pufels (Southern

wever, is more or less the same. Time scale as in Fig. 5.


soni (Solien), a guide form for the lower Spathian.

Setting the base of the Olenekian immediately above

the N. dieneri–N. cristagalli fauna, the y13C-peak in

Payne et al. (2004) will locate in the basal Smithian,

correlative with our positive carbon isotope excursion

at Pufels (Fig. 6). For Pufels, we set the base of the

Olenekian by the FAD of Pachycladina obliqua and

somewhat below the change from normal to reversed

magnetisation within the Seis Member (from Scholger

et al., 2000) that was re-dated by conodont (Farabegoli

and Perri, 1998, and own data) and magnetostrati-

graphic data (Tong et al., 2005). This places the base

about 3 m below the sample Bu 45 with the Smithian

Pachycladina obliqua Staesche, this sample being

close to the top of the only normal interval within the

Seis Member. The bulk of the positive y13C shift then

falls at Bu 48, about 1/6 up the Smithian.

In the Losar Tethyan section at Spiti in India, a 2xpositive carbon excursion with an apex at 3x, as at

Pufels, was reported by Atudorei (1999) near the base

of the lower Olenekian Flemingites Beds. Unfortunate-

ly, no ammonoid and conodont ranges were quoted for

this stratigraphic level and the recently published litho-,

sequence- and biostratigraphic revisions by Bhargava

et al. (2004) and Krystyn et al. (2004) suggest that the

excursion should have been assigned to immediately

below the base of the Smithian, thus preceding some-

what the one at Pufels. Nonetheless, this apparent

discrepancy is still within the uncertainties of biostra-

tigraphy. Atudorei (1999) describes a similar positive

excursion, up to 2.3x, also around the Gandarian–

Smithian boundary at Wadi Alwa in Oman.

The y13C values of the South China Chaohu sec-

tion (GSSP candidate for the Olenekian) show trends

similar to the above marine sections (Tong et al.,

2005), but the values in the Brahmanian and lower

Smithian are ~2x lower. Only in the uppermost

Gandarian and lowermost Smithian do the y13Cvalues reach ~1x level, with one negative value

reported at the base of the Olenekian. Note, however,

that this is a condensed section, with resolution much

lower than that at Pufels.

A positive y13C-shift occurs also in the lower cycle

11 of the Bernburg Formation within the lacustrine

deposits of the Germanic Basin (Korte and Kozur, in

press). In this section (borehole Halle-Sud), the base

of the Olenekian lies within cycle 7 (based on con-

chostracans), about 300 ka below the apex of the

positive shift in cycle 11 (Kozur and Bachmann,

2005; Korte and Kozur, in press).

The upper Smithian is characterised by low y13Cvalues, about �1x, in our section (Fig. 6), �1 to

�3x for the three sections of South China (Payne et

al., 2004), �3x for the Losar section near Spiti

(Atudorei, 1999) and much of the Smithian at Chaohu

(Tong et al., 2005). In Wadi Alwa, Oman, the y13Cvalues are mostly N1x, but they dip to ~0x in the

middle and upper Smithian (Atudorei, 1999).

For the early to middle Spathian no data were gen-

erated in the present study, but the gap is covered by

literature data, with y13C values at ~2.5 to 4x around

the Smithian–Spathian boundary and 0 to�2x higher

up in the Spathian (Atudorei, 1999; Payne et al., 2004;

Tong et al., 2005). This is followed by our data from

Palazzo Adriano, Sicily (Fig. 6) that show a rise from

0.3x to about 4x in the upper Spathian Triasso-

spathodus sosioensis and Chiosella gondolelloides

Zones and a subsequent rapid drop of nearly 3x in

the lower Aegean Chiosella timorensis Zone (Appen-

dix A). A positive y13C excursion of similar amplitude

with a maximum somewhat above the base of the

Anisian was reported by Payne et al. (2004) from

China, but taking the conodont ranges into account,

this excursion can be at the same stratigraphic level as

in the Palazzo Adriano. A potentially coeval y13Cincrease was observed also in the Niti Limestone

Member and the lower Himalayan Muschelkalk Mem-

ber (Mikim Fm.), within the Losar section at Spiti

(Atudorei, 1999), with a maximum in the lowermost

Anisian C. timorensis Zone (Krystyn et al., 2004).

Such high uppermost Spathian to lower Aegean (C.

timorensis Zone) values, 4 to 5x, were recorded

furthermore in the Desli Caira section (Dobrogea,

Romania) and the Kciras section of Albania (Atu-

dorei, 1999). In the Albanian section, however, C.

gondolelloides (Bender) was misidentified as C.

timorensis (Nogami) and the base of the Anisian

was likely drawn somewhat too deep. This positive

excursion was not detected in the Chaohu section

(Tong et al., 2005), but the position of the Olene-

kian–Anisian boundary in this area is not well known.

The Middle Triassic brachiopod and whole rock

y13C data remain low (Fig. 5), with a considerable

scatter, between �1 and 2x, but no clear excursions.

Only slightly more positive y13C values, ~2x, were

reported by Atudorei (1999) for a number of Tethyan


localities and by Payne et al. (2004) for the Chinese

sections.

The Late Triassic trend commences with an overall

rise of 3x in the Carnian (Fig. 5). Superimposed on

this rise is a short negative drop, of ~1.5x, in the late

Cordevolian and Julian (Fig. 7). Subsequently, in the

Early Norian, the y13C remains high, between 3 and

4x, dropping by ~1x in the Middle Norian. Note,

‰ (

V-P

DB

)

-4

-3

-2

-1

0

1

2

3

4

δ 13

C

δ 18

O

Carnian

Cordevolian Julian

aon aonoides

diebeli-noah tethydis-noah n. carp.

Tuval.

canad.

St. Cassian, brachiopods (Mn < 250 ppm; Sr > 400 ppm)St. Cassian, brachiopods (Mn > 250 ppm; Sr < 400 ppm)St. Cassian, aragonitic fossilKöveskál, whole rockPalazzo Adriano, whole rockPula Sándorhegy, brachiopod (Mn > 250 ppm)Silická Brezová, brachiopods(Mn < 250 ppm; Sr > 400 ppm)

Silická Brezová, whole rock

Silická Brezová, brachiopods(Mn > 250 ppm; Sr < 400 ppm)

austriacum dill. welleri

Fig. 7. Carbon and oxygen isotope variations in the Carnian. A

distinct increase in y18O can be seen from the late Cordevolian to

the early Julian, a trend established from well-preserved brachio-

pods from the St. Cassian.

there is a gap in the Middle Norian data because the

samples from Silicka Brezova that cover this time

interval could not be assigned definitive stratigraphic

positions due to complications of sedimentology and

stratigraphy. However, Gawlick and Bohm (2000)

studied a comparable time interval in the distal peri-

platform of the Northern Alps and observed a gradual

decline in y13C, from ~4 to 2.5x, similar to the

Silicka Brezova data at 2.5x, and Muttoni et al.

(2004) reported declining values, from 2.5 to b1x,

in the late Middle Norian for the Pizzo Mondello

section (Sicily).

In the Rhaetian, brachiopod and whole rock data

are ~1.9F0.7x. For the latest Triassic no data were

generated in the present study, but a strong negative

excursion at the Triassic–Jurassic boundary (Fig. 5),

with two pronounced negative peaks of up to �4x,

was documented already by Palfy et al. (2001) for

whole rock carbonates.

5.2. Oxygen isotopes

The y18O data for the brachiopod shells are sepa-

rated into their Tethyan and Muschelkalk provenance

(Fig. 8, Appendix A). Because oxygen isotope data

are prone to diagenetic alteration (Veizer, 1983a,b),

we exclude samples with Mn concentrations in excess

of 250 ppm and/or Sr contents of less than 400 ppm,

as well as the shells having alteration features visible

in SEM, from further consideration. For the Tethyan

realm, the whole rock y18O values are similar to those

of the well-preserved brachiopod shells. Only the

samples from the Pufels section have significantly

lower y18O values, up to �8x (Fig. 8), probably

due to diagenetic alteration.

Due to the absence of Early Triassic articulate

brachiopods in the investigated sections, two oxygen

isotope data from Korte et al. (2004a) for phosphatic

conodonts are included for comparison. Carbonate

and phosphate phases of modern carbonate shells, at

temperatures between 20 and 30 8C, are offset by

approximately 8.5x (Longinelli and Nuti, 1973; see

also Iacumin et al., 1996) and this was taken into

account in Fig. 8.

Overall, the y18O values for the well-preserved bra-

chiopods (Fig. 8) vary from �3.9 to �0.6x and�6.2

to �2.0x for the Tethys and the Muschelkalk Sea,

respectively. The spread in the data is discussed below.

Fig. 8. y18O values of Triassic LMC brachiopods, three aragonitic samples, and whole rock carbonates. The data sets are subdivided by

depositional regions (Tethys and Muschelkalk Sea) and trace element (Mn, Sr) concentrations; the latter as in Fig. 5. In addition, four

brachiopod shells are weathered or show alteration features in SEM (MK/base Pelsonian, MK/cycloides-Bank g at Schwieberdingen, SB 23,

Fels 21). The oxygen isotope values of these samples are likely altered. Time scale as in Fig. 5.


The whole rock y18O data for the Late Carnian and

Norian from the Silicka Brezova section are relatively

high, up to �0.3x, if compared to the articulate

brachiopod data from the adjacent time spans (Early

Carnian and Rhaetian). Note nevertheless that similar,

�1.2 to +0.1x, whole rock values were observed

also for the distal periplatform carbonates of the

Northern Alps by Gawlick and Bohm (2000).

5.3. d13C and d18O scatter of coeval brachiopods

The y13C and y18O values have a spread of more

than 2x for even the well-preserved coeval brachio-

pod shells (Figs. 5, 7 and 8). To some extent, this

scatter is due to the compressed time axis, which does

not permit resolution of short-term variations in the

isotope signal. Nevertheless, scatter in this range is

the norm rather than an exception for ancient

(Bruckschen and Veizer, 1997; Veizer et al., 1999)

as well as modern brachiopods (Brand et al., 2003).

For oxygen isotopes, the most frequent cause is the

variability in ambient seawater temperature, due to

bathymetric differences or seasonal upwelling (James

et al., 1997). Carbon isotopic variability within and

between brachiopod shells is also a common obser-

vation (e.g. Popp et al., 1986; Veizer et al., 1986;

Grossman et al., 1991, 1993) in both ancient and

modern specimens, as is the covariance of y13C and

y18O (Veizer et al., 1999; Brand et al., 2003). Such

variability in y13C is often due to incorporation of

respiratory CO2 during shell secretion, or may reflect

habitat and microhabitat factors (Curry and Fallick,

2002). Seasonal or taxonomic phenomena also play a

role (Brand et al., 2003). In an extreme case, devia-

tions of several permil in y13C and y18O were ob-

served within a single brachiopod shell from

Washington State (Auclair et al., 2003), but this

sample was collected from an intertidal zone. All

the above qualifications notwithstanding, the usual

deviations from isotopic equilibrium for brachiopods

are mostly within the F1x range (Carpenter and

Lohmann, 1995).


6. Discussion

6.1. Carbon isotopes

Model seawater y13C variations are usually as-

sumed to be a response to burial and re-oxidation of12C enriched organic matter within the ocean–atmo-

sphere system, caused by a plethora of factors, such as

CO2 levels, nutrient supply, rate of sedimentation, net

primary productivity, biological isotope fractionation,

or sea level changes (Scholle and Arthur, 1980;

Zachos et al., 1989; Holser, 1997; Jenkyns, 1996;

Hayes et al., 1999; Kump and Arthur, 1999). Other

factors may include input of volcanic CO2 into the

ocean–atmosphere system, which in the case of the

Siberian trap volcanism (Renne et al., 1995) may have

contributed ~1 to 1.5x to the large negative carbon

isotope shift at the Permian–Triassic boundary (Ber-

ner, 2002). Short-term negative y13CDIC excursions

may also be produced by sudden methane bursts

(Dickens et al., 1997) or by an oceanic overturn, the

latter resulting in the upwelling of 12C enriched (an-

oxic) bottom waters (Knoll et al., 1996). This last

scenario is assumed by us (Korte et al., 2004a) to

have been a major contributor to the 4x negative

y13C shift at the Permian–Triassic boundary.

Considering the open marine character of the in-

vestigated sediments (except for parts of the

Muschelkalk) and the coherence of overlapping sec-

tions, we believe that the previously described y13Cfluctuations reflect real secular variations of global, or

at least regional, significance.

Data from the present study and literature (Atu-

dorei, 1999; Korte et al., 2004a; Payne et al., 2004;

Korte and Kozur, 2005; Tong et al., 2005) suggest that

the Early to lowermost Middle Triassic Tethyan sec-

tions have similar carbon isotope fluctuations. This

despite differences in facies and bathymetry. They can

therefore potentially serve as chemostratigraphic mar-

kers for the Tethyan Realm or beyond. We suggest

that the following carbon isotope excursions may be

important stratigraphical markers:

(1) The pronounced negative excursion within the

uppermost Permian and lowermost Triassic that

culminated at the base of the Triassic (base of

Hindeodus parvus Zone) and in the lower Isar-

cicella isarcica Zone. This geochemical event is

discussed in several publications (e.g. Holser

and Magaritz, 1987; Baud et al., 1989; Holser

et al., 1989; Jin et al., 2000) and in detail in

Korte et al. (2004a,b,c) and Korte and Kozur

(2005).

(2) The positive y13C excursion in the late Gange-

tian reported by Korte et al. (2004a) for the

Abadeh section in Iran, by Payne et al. (2004)

for China, and in this study for the Palazzo

Adriano section in Sicily.

(3) A short positive excursion, of 2 to 3x mag-

nitude, in the basal Smithian reported here as

well as in Atudorei (1999), Tong et al. (2005)

and Korte and Kozur (in press). Much higher

values, up to 8x, were reported by Payne et

al. (2004).

(4) A negative excursion in the middle and upper

Smithian, with y13C values between �2 and

�3x, reported by Atudorei (1999) for the

Spiti section in India and by Payne et al.

(2004) and Tong et al. (2005) for the Olenekian

GSSP candidate section at Chaohu in South

China. At Pufels, two negative drops are present,

with the lowest values in the upper Smithian.

(5) A distinct positive excursion in the uppermost

Smithian and lower Spathian, with y13C values

up to 4x according to Atudorei (1999) and

Tong et al. (2005), or up to 2.5x according to

Payne et al. (2004).

(6) Low isotope values, 0 to �2x, in the lower part

of the upper Spathian (Atudorei, 1999; Payne et

al., 2004; Tong et al., 2005).

(7) A pronounced positive excursion in the upper-

most Olenekian and lowermost Anisian, with a

maximum at 4 to 5x (present study; Atudorei,

1999; Atudorei et al., 2002; Payne et al., 2004).

The strong shifts in the Early Triassic seawater

y13C likely reflect global instability of the carbon

cycle, rather than any local, diagenetic or facies-spe-

cific effects. Taking into consideration that the Lower

Triassic had a short duration of ~5.6 Ma (Kozur,

2003a,b), these shifts during times of limited biolog-

ical recovery (Payne et al., 2004) from the massive

end-Permian extinction event must have been of rel-

atively short duration.

It is difficult to assign specific causations to any

single carbon isotope excursion, but two shifts may


tentatively be attributed to changes in bioproductivity.

The positive excursion at the Smithian–Spathian

boundary (Atudorei, 1999; Payne et al., 2004; Tong

et al., 2005) is contemporaneous with the termination

of the radiolarite gap in the Tethys and in the low

latitude Panthalassa, marking the recovery of marine

biota, particularly of siliceous plankton (Kozur,

1998a,b). The large and relatively short-lived peak

at the Olenekian–Anisian boundary (Fig. 6) coincides

with the recovery of land plants, as documented by the

palynoflora in the central and western Tethys and in

the Germanic Basin (Kozur, 1999). Here, the rich but

monotonous Densoisporites neijburgi (miospore of

Pleuromeia) association (e.g. the upper Detfurth and

Hardegsen Formations of the Germanic Basin, the

Dorikranites and Tirolites cassianus beds of the Pri-

caspian Depression and Mangyshlak, the lower Cso-

pak Marl Formation of Hungary) is replaced by the

very rich and diverse Voltziaceaesporites heteromor-

phus sporomorph association. The latter is ubiquitous

in the Solling Formation of the Germanic Basin as

well as in the uppermost Csopak Marl and the lower-

most AszofI Dolomite Formations of Hungary. This

sudden increase in land plant abundance and diversity

may have withdrawn light 12C from the atmosphere/

ocean-system, leading to the observed y13C rise. Nev-

ertheless, it is doubtful if this alone could have gen-

erated a nearly 4x shift.

During the Anisian, the recovery of the land plants

continued, yet the carbon isotope values are low and

remain low until the Middle–Late Triassic boundary

(Fig. 5). The higher scatter and the somewhat lower

values of our Muschelkalk brachiopod shells, if com-

pared to the literature whole rock data, will be dis-

cussed below together with their oxygen isotope data.

The Middle Triassic time interval is characterised by a

widespread sea level rise, with transgressions along the

entire Tethyanmargins, in the Germanic Basin (Rot and

Muschelkalk), in western North America and in the

Pricaspian depression. It was accompanied by a wide-

spread deposition of organic rich sediments. Examples

are the AnisianGuttenstein Limestone and the bcalcairevermiculaireQ (bWurmlikalkeQ) in wide parts of the

Tethys, the Anisian–Ladinian bGrenzbitumenzoneQ ofSwitzerland, the contemporaneous Pestis Shale of the

northern Apuseni Mts. in Romania, and the Ladinian

Partnach Shales in the Northern Alps. The convention-

al interpretation would demand that the times of rising

sea levels and increased deposition of sediments rich in

organic matter should correspond to heavier carbon

isotope values (Scholle and Arthur, 1980; Weissert

and Lini, 1991; Follmi et al., 1994; Jenkyns, 1996),

but this is at odds with the observed trend (Fig. 5). In

addition, and in contrast to shelf seas, the open marine

pelagic domain remained oxic, as attested by red radi-

olarites of the Tethys and Panthalassa (Kozur and

Mock, 1988; Pillevuit, 1993; Imoto and Kozur,

1997). Nor does evidence for a large-scale deposition

of organic facies exist on land, since coal deposits were

sparse until the uppermost Ladinian (Retallack et al.,

1996). These contradictory observations make it diffi-

cult to advocate a specific scenario for the low Middle

Triassic y13C values. This time interval, nevertheless,

coincides with a widespread explosive felsic to inter-

mediate volcanism across large sections of the Tethys

(Fig. 5) and volcanic CO2 may have contributed to the

low Middle Triassic y13C values.

The re-emergence of coal swamps and peatlands

may have been the cause of the 3.5x rise in the

Carnian (Fig. 5) and of the predominately 13C-

enriched values thereafter, but note that the rise in

y13C commences somewhat later than the renewal of

coal swamp sedimentation. In addition, a N1.5x drop

is indicated for the middle Carnian (Fig. 6), following

a short-term sea-level drop (Haq et al., 1987). This

drop is coincident with a short phase of strong mafic,

intermediate and felsic volcanism across large por-

tions of the Tethys and with the widespread deposition

of black anoxic or dysaerobic shales.

We can only speculate about the causes of some

1x decline in carbon isotope values within the Mid-

dle Norian because little is known about this time

span. In the Tethys, the Hallstatter Limestone (Hall-

statt, Silicka Brezova) and the Hauptdolomit facies are

similar throughout the entire Carnian and Norian, but

facies changes are documented in the Newark Basin

of North America, where black shales with some

bituminous limestone of the Early Norian Lockatong

Formation (Channell et al., 2003; Muttoni et al., 2004)

are overlain by the predominantly red sediments of the

lower Passaic Formation. Contemporaneous deposits

of the northern arid Tethyan Belt and the Germanic

Basin are devoid of coals and organic-rich beds. Even

sporomorphs are very rare. Decreased organic carbon

burial rates could therefore be one possible scenario

for the decline in the y13C values.


The decline in carbon isotope values from the

upper Norian to the Rhaetian (Fig. 5) is coeval with

abundant coal deposits and plant detritus on conti-

nents, bituminous limestones in marginal seas (Kos-

sen Limestone), and dark limestones and marls rich in

organic material in the open sea (Zlambach Beds). In

general, the Rhaetian sediments have a higher detrital

component (Zlambach Marls above Hallstatt Lime-

stones, Kossen Beds above Hauptdolomit and Platten-

kalk in marine beds, Rhatian sandstone above Norian

Steinmergel in continental and partly also in the shal-

low marine Germanic Basin). All this indicates higher

relief due to tectonic unrest and more humid climate.

On continents, sediments are rich in plant remains and

sporomorphs (Germanic Basin) even in the former

low latitude dry girdle, indicating more humid condi-

tions (Hussner et al., 1996). In view of these observa-

tions, the observed y13C decline is opposite to that

expected from geological considerations. It may be of

some interest that eruptions of large plateau basalts in

the Central Atlantic Magmatic Province (CAMP)

commenced in the Rhaetian, but mostly in its later

part, culminating at the Rhaetian–Liassic boundary.

Note also that the Triassic–Jurassic extinction event

had already begun at the base of the Rhaetian.

6.2. Oxygen isotopes

The y18O of brachiopods (Fig. 8) is affected by

three variables, the temperature, the y18O, and poten-

tially also the pH of the Triassic seawater. The St.

Cassian (early to middle Carnian) and the Kossen

Limestone (Rhaetian) brachiopods, with y18O be-

tween �0.6 and �3.4x V-PDB, coexisted with

reef-building corals. The latter require permanent

water temperatures of ~18 to 34 8C, although short-

term drops down to 16 8C can be tolerated (Kleypas et

al., 1999). However, perpetual temperatures below 18

8C, or repeated short-term drops below 16 8C, arelethal. The Tethyan brachiopod samples are either

from tropical open marine environments or from

large intraplatformal basins (Kossen Beds) and one

can therefore assume that the effects of evaporation,

dilution or pH were not extensive. Assuming y18O of

0x (V-SMOW) for Triassic seawater, and applying

the equations of O’Neil et al. (1969) and Hays and

Grossman (1991), the �0.6 and �3.4x y18O range

of brachiopods translates into temperatures of ~18 to

32 8C, in general agreement with the tolerance limits

of reef-building corals.

For the Germanic Muschelkalk Sea brachiopods

(y18O of �2.0 to �6.2x) the situation is more com-

plex. Assuming again 0x V-SMOW for the

Muschelkalk seawater, the calculated temperature

range is from 25 to 47.5 8C, in excess of the ~38 8Ctolerance limit for higher organisms (Brock, 1985;

Rothschild and Mancinelli, 2001). This discrepancy

can be resolved in two ways. First, the three most

depleted samples – MK/Meso 32 (�6.23x), MK/

Meso 34 (�4.75x), MK/Meso 41 (�5.59x) – al-

though characterised as well preserved, may be reset

after all. If we reject these three samples, the y18Orange of the Muschelkalk brachiopods only varies

from �2.0 to �4.4x. The corresponding calculated

water temperatures would be from 25 to 37 8C. Theseare marginally acceptable values for an epi-continental

sea in arid latitudes at approximately 258 to 358N. Asecond, and more likely, explanation is that the

Muschelkalk Sea of the Germanic Basin was relatively

restricted, although a connection to the Tethys existed

for the most part. As a consequence, it was subject to a

strong influx of meteoric waters, as attested by the fact

that the basin becomes increasingly brackish up-sec-

tion in the north and east (Kozur, 1976). This scenario

is supported also by 87Sr/86Sr values that are higher

than those for the coeval open sea seawater (Korte et

al., 2003). The influx of meteoric water would have

lowered the y18O of Muschelkalk seawater to about

�2x (V-SMOW) or less, in which case even the

lowest y18O values of �6.2x (V-PDB) could still

yield a tolerable temperature of 36.5 8C. Tentatively,this scenario could also be the reason for the large

scatter and relatively low y13C values of the

Muschelkalk brachiopods, if compared to the coeval

whole rock data of Atudorei (1999) for the Tethys and

of Payne et al. (2004) for the Chinese successions.

In the Early Carnian (Fig. 7), the oxygen isotope

values rise by about 2x within the Cordevolian, from

�3x in the early/middle aon Zone to �1x in the

late aon/early aonoides Zone, and decline again to

�3x in the aonoides Zone (but note that only two

values exist for the late aonoides Zone). These South

Alpine late Cordevolian to lowermost Julian fluctua-

tions would suggest a rapid temperature decline of

almost 10 8C, to temperatures around 20 8C. For thesame South Alpine region, Mutti and Weissert (1995)


proposed a salinity increase for ambient seawater.

However, a y18O shift of about 2x would require a

relatively large increase in salinity of ~6x (Craig,

1966). Thus a combination of a moderate temperature

drop with a moderate increase in seawater salinity

may be the most likely explanation.

The y18O of the late Carnian Silicka Brezova bra-

chiopods are not interpreted in this study because only

4 brachiopods were characterised as well preserved by

trace element criteria. Even for these, the one sample

that was studied by SEM (Fig. 3f) showed some

recrystallisation features. Four y18O values of well-

preserved brachiopods of the Anisian FelsIors sec-

tions that vary between �1 and �3.9x indicate

temperatures between 20 and 34.5 8C.The overall similarity of the y18O values for the

coeval well-preserved brachiopod shells and whole

rock carbonates (Fig. 8) is somewhat unexpected.

This observation suggests that the lithification of car-

bonate sediments must have been achieved rapidly

during early diagenesis, and was completed while

the pore waters were still in at least a partial diffusive

interchange with the overlying seawater (see also the

discussion in Veizer et al., 1999). As already men-

tioned, the 18O depleted whole rock carbonates of the

Pufels section reflect a clear meteoric influence during

diagenesis. On the other hand, the y18O values of the

Late Carnian and Norian whole rock carbonates are

high if compared to the bracketing Carnian and Rhae-

tian brachiopod data. However, these carbonates, the

locally chert-bearing Hallstatt Limestones at Silicka

Brezova (Slovakia), were deposited on the slope of

the Meliata Ocean and they contain an assemblage of

palaeopsychrosphaeric ostracods (sensu Kozur, 1991)

and ostracods that prefer shallow water. Faunas with a

high percentage of palaeopsychrosphaeric ostracods

are characteristic of water depth of about 150–200 m

(Kozur, 1998c), close to the thermocline, that even in

tropical habitats is considerably cooler than the shal-

low water. Such cooler temperature during early dia-

genesis may have been the reason for the relatively

heavy y18O values of the whole rock Norian carbo-

nates. As pointed out above, the Rhaetian brachio-

pods, collected from the Kossen Beds, are ~1xdepleted in 18O relative to the Norian whole rock

carbonates. The Kossen sediments were deposited in

an intraplatformal basin that was shielded from cold

bottom currents due to its shallow water depths of 20–

80 m (Urlichs, 1972). This resulted in distinctly higher

temperatures than those prevalent on the ocean slope

at Silicka Brezova, with an estimated depth of ~150 to

200 m. Hence the oxygen isotopic shift at the Norian–

Rhaetian transition (Fig. 8) may not be a reflection of

climate warming.

7. Conclusions

We have analysed 318 samples of articulate bra-

chiopod shells and whole rock carbonates for carbon

and oxygen isotopes, thus generating the first baseline

trends for y13C and y18O of Triassic seawater. The

Early Triassic and earliest Middle Triassic data indicate

the presence of about seven short-lived y13C excur-

sions. Those at the Smithian–Spathian and Olenekian–

Anisian boundaries may be tentatively attributed to

changes in bioproductivity, but the causes for the

other excursions are, at this stage, equivocal. The

low y13C values during the Middle Triassic coincide

with a strong explosive felsic to intermediate volca-

nism in large portions of the Tethys and volcanism may

be partly responsible for the observed 13C depleted

values. The 3x rise in y13C values during the Carnian

coincides with the re-establishment of large-scale coal

deposition that may have resulted in withdrawal of

light 12C from the atmosphere/hydrosphere system.

Assuming y18O of 0x (V-SMOW) for Triassic

seawater, the oxygen isotope values for the Late Trias-

sic brachiopods (�0.6 and�3.4xV-PDB) yield water

temperatures of ~18 to 32 8C, in rough agreement with

the temperature tolerance of the coexisting reef-build-

ing corals. The 2x increase in y18O, from�3 to�1x,

in the early Carnian likely reflects a combined impact

of temperature decline and increase in seawater salinity

(see Mutti and Weissert, 1995). For the Muschelkalk,

the y18O values of the well-preserved brachiopods

range from �2 to �6.2x, likely a reflection of an

influx of meteoric waters into the arid Germanic Basin

that resulted in lowered y18O of ambient waters.

Acknowledgements

This project was financially supported by the

Deutsche Forschungsgemeinschaft (Leibniz-Prize, Ve

112/8-1; grant Ve 112/12-1) and by the Deutsche


Akademie der Naturforscher Leopoldina (BMBF-LPD

9901/838). Excavation of the Silicka Brezova section

was financially supported by the US National Science

Foundation (NSF) grant (EAR 94-17895) to J. Chan-

nell and was carried out by the late R. Mock. Addi-

tional samples were contributed by H. Hagdorn, M.

Urlichs and J. Michalık. The analytical support of H.

Strauss (at Ruhr-Universitat Bochum), C. Spotl (at

Universitat Innsbruck) and P. Bruckschen (at Texas

A&M University, College Station) is appreciated. We

thank H. Jenkyns, P. Swart and H. Weissert for

reviews of the manuscript and helpful annotations.

Appendix A. Supplementary material

Supplementary data associated with this article can

be found, in the online version, at doi:10.1016/j.

palaeo.2005.05.018.

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δ 13C and δ 18O values of Triassic brachiopods and carbonate rocks as proxies for coeval seawater...

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