Int J Earth Sci (Geol Rundsch) (2009) 98:677–695
DOI 10.1007/s00531-007-0264-1
ORIGINAL PAPER
Authigenic carbonate precipitates from the NE Black Sea:
a mineralogical, geochemical, and lipid biomarker study
A. Bahr Æ T. Pape Æ G. Bohrmann Æ A. Mazzini Æ
M. Haeckel Æ A. Reitz Æ M. Ivanov
Received: 6 June 2007 / Accepted: 6 October 2007 / Published online: 31 October 2007
Ó Springer-Verlag 2007
Abstract Carbonate precipitates recovered from 2,000 m
water depth at the Dolgovskoy Mound (Shatsky Ridge, north
eastern Black Sea) were studied using mineralogical, geochemical and lipid biomarker analyses. The carbonates
differ in shape from simple pavements to cavernous structures with thick microbial mats attached to their lower side
and within cavities. Low d13C values measured on carbonates (-41 to -32% V-PDB) and extracted lipid biomarkers
indicate that anaerobic oxidation of methane (AOM) played
a crucial role in precipitating these carbonates. The internal
structure of the carbonates is dominated by finely laminated
coccolith ooze and homogeneous clay layers, both cemented
by micritic high-magnesium calcite (HMC), and pure, botryoidal, yellowish low-magnesium calcite (LMC) grown in
direct contact to microbial mats. d18O measurements suggest that the authigenic HMC precipitated in equilibrium
with the Black Sea bottom water while the yellowish LMC
rims have been growing in slightly 18O-depleted interstitial
water. Although precipitated under significantly different
A. Bahr (&) T. Pape G. Bohrmann
MARUM-Zentrum für Marine Umweltwissenschaften
der Universität Bremen, PO Box 330440,
Bremen 28334, Germany
e-mail: bahr@uni-bremen.de
A. Mazzini
Physics of Geological Processes (PGP),
University of Oslo, PO Box 1048, Blindern, Norway
M. Haeckel A. Reitz
IFM-GEOMAR, Wischhofstr. 1-3, Kiel 24148, Germany
M. Ivanov
UNESCO MSU Center for Marine Geology and Geophysics,
Faculty of Geology, Moscow State University,
Vorobjevy Gory, Moscow 119899, Russia
environmental conditions, especially with respect to methane availability, all analysed carbonate samples show lipid
patterns that are typical for ANME-1 dominated AOM
consortia, in the case of the HMC samples with significant
contributions of allochthonous components of marine and
terrestrial origin, reflecting the hemipelagic nature of the
primary sediment.
Keywords Black Sea AOM Authigenic carbonates Methane seepage Dolgovskoy Mound
Introduction
Studies on authigenic carbonates sampled from various
geological settings as Hydrate Ridge (e.g., Bohrmann et al.
1998; Greinert et al. 2001; Teichert et al. 2005), Blake
Ridge (e.g., Naehr et al. 2000; Pierre et al. 2000; Rodriguez
et al. 2000) or the Black Sea (e.g., Mazzini et al. 2004;
Michaelis et al. 2002; Peckmann et al. 2001; Pape et al.
2005; Reitner et al. 2005a; Stadnitskaia et al. 2005) have
shown that their precipitation is a common process related
to methane seepage and concomitant microbial consumption of methane. In anoxic and sulphate-rich environments,
methane is consumed by the anaerobic oxidation of
methane (AOM), performed by a consortium of methanotrophic archaea and sulphate reducing proteobacteria
(Boetius et al. 2000) according to the net reaction:
CH4 þ SO2
4 ! HCO3 þ HS þ H2 O:
ð1Þ
As a result of AOM bicarbonate is produced, increasing the
alkalinity of the interstitial water, which leads to carbonate
precipitation. The study of lipids incorporated in authigenic
carbonates provides additional insight into the phylogenetic
affiliation of AOM communities indirectly inducing the
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678
carbonate formation. Until now, three phylogenetic distinct
clusters of anaerobic methanotrophic archaea, called
ANME-1 and ANME-2, both known from the Black Sea,
and ANME-3, so far not observed in the Black Sea (Knittel
et al. 2005) are described. Based on studies of carbonate
precipitates from ca. 230 m water depth in the northwestern
Black Sea, Reitner et al. (2005b) suggested that the composition of the AOM consortia affects indirectly the
carbonate mineralogy in the way that different groups of the
AOM-involved archaea have different methane turnover
rates, low for ANME-1, and high for ANME-2 (Nauhaus
et al. 2005); accordingly, a high methane turnover rate
reduces the sulphate concentration in the pore water faster
than a low AOM rate. The inhibition of the Mg incorporation into calcite at the presence of sulphate accompanied by
a distortion of the crystal lattice (Kralj et al. 2004), as well
as the preferred precipitation of aragonite over calcite at
high sulphate concentrations has been known for some time
(see, e.g., Burton 1993). Hence, ANME-1 dominated
assemblages were found associated with aragonite sections
and ANME-2 assemblages at HMC-dominated parts of the
carbonate structures (Reitner et al. 2005b).
A suite of 13C-depleted biomarker lipids is typically
synthesized by AOM-performing communities. For
instance, isoprenoidal glycerol diether lipids like archaeol
and more functionalised derivatives thereof are diagnostic
for archaea (Elvert et al. 1999; Thiel et al. 1999; Hinrichs
et al. 2000), while carboxylic acids, non-isoprenoidal
glycerol monoether (monoalkyl glycerol monoether,
MAGE) and diether (dialkyl glycerol diether, DAGE) lipids
are considered to be produced by bacteria (e.g., Pancost
et al. 2001; Thiel et al. 2001a). Although the source
organisms of non-isoprenoidal diether lipids eluded cultivation so far and, thus, are still uncertain, several hints exist
that these compounds originate from sulphate-reducing dproteobacteria (e.g., Pancost and Sinninghe Damsté 2003).
Based on characteristic lipid distribution patterns found for
phylogenetically distinct anaerobic methanotrophic consortia at several widely distributed sites, Blumenberg et al.
(2004) established indices to evaluate the predominant
AOM populations in recent and fossil samples.
In the literature much attention has been paid on carbonate build-ups found on the northwestern Black Sea shelf
and upper slope (e.g., Michaelis et al. 2002; Pape et al.
2005; Peckmann et al. 2001; Reitner et al. 2005b). However, only few data exist on the more laterally extensive
carbonate pavements found elsewhere in the deep Black
Sea (Mazzini et al. 2004; Pape et al. 2005; Stadnitskaia
et al. 2005). In this paper we present data obtained from
carbonates retrieved from the Dolgovskoy Mound (DM) in
the deep (i.e. *2,000 m water depth) north eastern Black
Sea. Using a combined methodological approach we give a
detailed assessment of the mineralogy and geochemistry of
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Int J Earth Sci (Geol Rundsch) (2009) 98:677–695
these precipitates as well as of the composition of associated AOM communities. The results are compared to
findings from other areas of the Black Sea.
Study area
The Black Sea originated during Cretaceous times as a backarc basin and experienced a complex history of changing
extensional and compressional tectonic regimes since then
(Nikishin et al. 2003). Due to a stable stratification with
freshened surface water (about 18% salinity) and more
saline water below ca. 160 m (22.5% salinity), derived from
the Mediterranean Sea (Öszoy and Ünlüata 1997), the contemporary Black Sea represents the world’s largest
permanently anoxic basin, with high methane concentrations
of up to 14 lM (Bohrmann et al. 2003) in the deep waters.
The only connection to the global ocean is provided by the
narrow and shallow Bosphorus Strait (sill depth: -35 m).
The typical sedimentary succession of late Pleistocene
to Holocene sediments in the basinal Black Sea comprises
laminated coccolith ooze (termed ‘‘Unit 1’’ after Ross and
Degens 1974) as the uppermost layer, underlain by a
laminated sapropel (‘‘Unit 2’’). Both are marine units and
have a combined thickness of ca. 50 cm. Below Unit 2,
several tens of meters of clayey-silty lacustrine mud (‘‘Unit
3’’) are found, representing the last glacial stage when the
Black Sea was disconnected from the global ocean due to
the low global sea level. The marine inflow occurred
around 9.4 kyr BP (Major et al. 2006), and the first sapropel-layers, indicating anoxia, were deposited at ca.
8 kyr BP (Lamy et al. 2006).
The DM is located east of the Crimean Peninsula on the
northwestern Shatsky Ridge, a structure of debated origin,
but possibly being a part of variscan deformed Precambrian
basement (Nikishin et al. 2003) overlain by Mesozoic
sediments (Afanasenkov et al. 2005). In the adjacent
Sorokin Trough a number of mud volcanoes, which were
formed due to the diapiric rise of clay-rich deposits from
the Oligocene—Lower Miocene Maikopian Formation (Ivanov et al. 1996), have been discovered. However, on the
Shatsky Ridge, no evidence for mud volcanic activity has
been observed so far.
The DM itself is a subcircular structure, ca. 800 m wide
and 70 m high. The sidescan sonar track taken during
training through research (TTR) 15-cruise in 2005 (Fig. 1)
shows irregular backscatter patterns with some strong
reflections at the centre of the DM indicating a complex
build-up of the sea floor (Fig. 1). Although no gas seepage
has been observed during the site survey with the TV-grab,
the abundance of AOM-related archaea and bacteria
forming thick mats give firm evidence for an elevated
methane concentration in the interstitial water at the DM.
Int J Earth Sci (Geol Rundsch) (2009) 98:677–695
679
Fig. 1 Location of the
carbonate sampling site
BS346GR during the TTR 15
cruise (R/V Professor
Logachev, June 2005) on the
Dolgovskoy Mound. The circle
in the map indicates the area
investigated as shown in the
detailed sidescan image. The
sidescan sonar line MAK58
over Dolgovskoy Mound and
the corresponding subbottom
profiler line between points A
and B (location see sidescan
image) were likewise obtained
during the TTR 15 cruise
Characteristic temperature and d18Owater values from the
abyssal Black Sea at the depth of our sampling site are 9°C
(Öszoy and Ünlüata 1997) and -1.77% (SMOW) (Rank
et al. 1999), respectively.
Sample description
Based on sidescan sonar and sub-bottom profiler data, a
TV-guided grab was deployed on top of the DM (station
BS346GR; 44°01.130 N, 36°41.400 E, 2,004 m water depth;
location see also Fig. 1) during TTR cruise 15 (June 2005)
onboard R/V Professor Logachev (Akhmetzhanov et al.
2007). The carbonate precipitates retrieved showed a
variety of forms: Some are slab-shaped calcite-cemented
hemipelagic sediments, representing carbonate Type-U1
according to the classification of Mazzini et al. (2004).
Other carbonates form complex structures (termed ‘‘complex type’’ in the following), up to ca. 50 cm in diameter
and with several centimetres thickness. These carbonates
were usually covered with carbonate cemented hemipelagic mud. The lower part of these carbonates is irregular,
often cavernous and coated by up to 2–3 cm thick microbial mats (Fig. 2) that also fill the numerous interconnected
cavities within the precipitates (Fig. 3). Two carbonate
specimens representing the complex type were chosen for
further analyses (Fig. 3).
Despite the apparent complex shape of the precipitates,
vertical sections cut through different carbonate pieces
highlight similarities. Both specimens shown in Fig. 4
(A-1, A-2 and B-1, B-2 are sub-samples of carbonate
pieces A and B, respectively) comprise micrite-cemented
coccolith ooze sediments (belonging to Unit 1) intercalated
by one or more micrite-cemented clay-rich layers (further
called ‘‘micritic layers’’), resembling Typ-MSa slabs from
Mazzini et al. (2004). In specimen A-1 one relatively soft
(less calcified) layer of this type is observed with a thickness of up to 2.8 cm while in specimen B-1 a succession of
three thin and indurated (well calcified) micritic layers is
alternated by cemented coccolith ooze layers (Fig. 4,
specimen B-1). The lower side of both specimens, A-1 and
B-1, is encrusted by botryoidal, mostly yellowish to reddish, sometimes whitish carbonates (in the further text
generally referred to as ‘‘yellowish calcite’’; Fig. 5a) that
also coat the internal part of the cavities (discerned as
‘‘cavity fill’’ in Fig. 4). Between the yellowish calcite and
the cemented coccolith ooze a transitional zone of dark
grey micritic carbonate developed in piece A-1 (‘‘TZ’’,
Fig. 4). This zone consists lithologically of heavily
cemented coccolith ooze laminations but shows a different
Fig. 2 Pink–orange microbial mat filling a cavity (the leftmost cavity
on Fig. 3b) on the lower side of the studied carbonate precipitate
retrieved at station BS346GR. The width of the photographed area is
ca. 2 cm
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Int J Earth Sci (Geol Rundsch) (2009) 98:677–695
Fig. 3 Photographs of the
sampled carbonate precipitates.
a Upper and b lower side of
precipitate A; c slightly oblique
view on upper and d lower side
of precipitate B. e Section
through piece A; the position of
the transect is marked as a
straight white line in A and B,
dashed white lines indicate the
transects shown in Fig. 4.
Arrows in b and d indicate the
openings of cavities on the
lower side of the precipitate.
The left most arrow in b
indicates the position of the
cavity shown in Fig. 2
carbonate mineralogy and geochemistry than the latter (see
‘‘Results’’). The spatial distribution of the above carbonate
microfacies in the studied samples is given in Fig. 4.
Scan CS 44) with an attached energy dispersive X-ray
spectrometer (EDX, instrument: EDAX 9800) at the Geoscience Department at the University of Bremen.
Methods
X-ray diffractometry
SEM
Ten samples (Table 1) of the slices A-1, B-1 and B-2 were
prepared for X-ray diffraction (XRD) analysis. Prior to
measurements, the samples were dried, ground in an agate
mortar and, after adding corundum (a-Al2O3) as an internal
After transecting the slabs, freshly broken samples were
analysed with a scanning electron microscope (SEM, Cam
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681
Fig. 4 Vertical transected parts
of specimen A (sub-samples
labelled A-1 and A-2) and B
(labelled B-1 and B-2).
Drawings show a schematic
distribution of the different
lithological facies that are
referred to in the text. White
circles indicate the positions
where the stable isotope
measurements were taken, the
shading in the interpretative
drawing to specimen A-1
specifies the parts that were
sampled for lipid biomarker
analyses. The three arrows in
specimen B-1 point at the
succession of three micritic
layers, each separated by a few
laminae of coccolith ooze.
‘‘Rim’’ indicates the more
calcified fringes of the micritic
layer in specimen A-1. The
samples are oriented with the
top facing upward
standard with the sample-to-standard ratio 1:5, homogenized with acetone and prepared as randomly oriented
powder slides. Measurements of samples A-1a to A-1e
were performed using an X’Pert PRO X-ray diffractometer
from PANalytical with an X-Celerator detector system at
the Geoscience Department at the University of Bremen.
The step size was set to 0.0167° covering a range between
3° and 85° 2h. The data were obtained with a fixed
divergence slit (0.38 mm), a Cu-Anode with kCuKa =
1.5406 and 45 kV source voltage. Samples B-1a to B-2a
were measured using a Philips SIETRONICS XRD SCAN
with an automated divergence slit and a Co-Anode
(kCoKa = 1.7903) with 40 kV covering a range of 18°–52°
2h. The visualisation and analyses of the measurements
was performed with the freeware program MacDiff 4.2.2.
provided by Rainer Petschick, Frankfurt University, Germany. All diffractograms were corrected for the positions
of the corundum peaks. The molar MgCO3-percentage of
calcite can be calculated by using the d(104) value of
calcite given in Å following the equation by Lumsden
(1979):
MgCO3 ðmol%Þ ¼ 100 ½333:33 dð104Þ 911:99:
ð2Þ
Stable carbon and oxygen isotope analyses
on carbonates
Thirty-seven samples for stable carbon and oxygen isotope
analyses were obtained from specimens A-1, A-2 and B-1
by using a Proxxon microdrill (for the exact position of the
samples cf. Fig. 4). The samples were dissolved in H3PO4
at 75°C and analysed with a Finnegan MAT 251 mass
spectrometer with an automated carbonate preparation
device at the Geoscience Department, University of Bremen. The ratios of bulk carbonate 18O/16O and 13C/12C are
given in the d-notation in % relative to the V-PDB standard. Analytical long-time standard deviation is ±0.07%
V-PDB for d18O and ±0.04% V-PDB for d13C.
Organic and inorganic carbon content, nitrogen content
For analyses of total organic carbon (TOC), carbonate and
nitrogen content, we measured six samples (Table 1) with
an Elementar Vario EL III. Each sample was analysed
untreated and after decalcification with 10% HCl solution
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Int J Earth Sci (Geol Rundsch) (2009) 98:677–695
Fig. 5 a Lower side of specimen A-1 in Fig. 4 showing the
botryoidal shape of the yellowish calcite; dark, rippled parts (arrow)
are coatings of clay minerals, pyrite, LMC, halite and organic matter;
b needle-shaped LMC crystals of the cavity fill (Fig. 4, specimen
B-1); c micritic HMC crystals (ML in Fig. 4, piece B-1); d SEM
photograph of a part of the coccolith ooze from specimen A-1 in
Fig. 4 containing abundant coccoliths derived from Emiliania huxleyi
(decalcification was done by adding drops of HCl to the
dried sample until no visible reaction occurred). The carbonate content is calculated from the difference of the total
carbon (TC) and TOC after decalcification:
of the dried sub-samples were repeatedly cleaned with
diluted pre-cleaned HCl and acetone. After crushing, the
carbonaceous material was covered with purified water and
dissolved by dropwise adding diluted HCl. The reaction
was stopped when about 10 wt.-% of the original material
were still left over. The reaction mixture (I) was decanted
and the remaining particles were repeatedly extracted with
methanol (CH3OH) and dichloromethane (CH2Cl2). Reaction mixture I and the organic extracts were combined in a
separatory funnel and extracted five times versus CH2Cl2
yielding the total organic extract (TOE). This TOE was
dried with a N2 stream and subjected to mild alkaline
hydrolysis using 6% KOH in CH3OH (2 h, 80°C) yielding
reaction mixture II. Reaction mixture II was repeatedly
extracted versus n-hexane to obtain the so-called neutral
lipid fraction comprising hydrocarbons and alcohols. Carboxylic acids were obtained by acidification (HCl, pH 1–2)
of the residual reaction mixture II and extraction with nhexane versus water. Hydrocarbons and alcohols in the
neutral lipid fraction were separated with a silica gel
(Merck silica gel 60) column by means of eluents of
increasing polarities.
For gas chromatographic analyses, alcohols and carboxylic acids were converted to their trimethylsilyl and
methylester derivatives, respectively, as previously
described (Pape et al. 2005).
Carbonate ðwt%Þ ¼ 8:333ðTC TOCÞ:
ð3Þ
Replicate measurements yielded an accuracy of
±0.75 wt% for carbonate content, ±0.09 wt% for TOC
and ±1.0 for TOC/N.
Lipid analyses on carbonate samples
From the three major lithological facies constructing the
carbonate precipitates (coccolith ooze, micritic layer, and
yellowish calcite) three fractions each were prepared
comprising of compounds extractable upon carbonate dissolution (fatty acids, hydrocarbons, and alcohols). All
fractions were analysed for concentration and stable carbon
isotopic composition of individual compounds.
Preparation and extraction
Lipids were prepared according to a modified method
previously described in Pape et al. (2005). About 25–30 g
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Dominated
by pyrite
683
Coccolith ooze
Micritic layer
Coccolith ooze
Yellowish calcite
Pyritic coating
Micritic layer
A-1e
B-1a
B-1b
B-1c
B-1d
B-2a
Weighted MgCO3 (mol%) = (IHMC 9 MgCO3,HMC + ILMC 9 MgCO3,LMC)/(IHMC + ILMC)
a
Intensities of corundum were measured at the (012) peak
I peak intensity, HMC high-Mg calcite, LMC low Mg-calcite, Cor corundum, TOC total organic carbon, C/N TOC/N, NA not analysed, NC not calculated
NA
NA
12.4
NC
9.21
2.9987
12.4
NC
NC
NC
NA
NA
NA
5.9
5.9
NC
NC
NC
3.64
3.0182
NC
NA
NA
NA
5.9
11.4
5.9
NC
NC
NC
3.0182
NC
11.4
3.0017
NC
NC
8.38
11.01
NC
NC
NA
0.62
96.8
NA
11.3
9.5
0.089
NC
0.3
NC
3.0352
NC
NC
0.97
10.3
11.3
3.0021
3.0052
10.83
7.07
0.39
90.2
11.5
0.034
0.6
3.0343
0.38
11.9
3.0003
For all fractions prepared, lipid biomarker distributions
were analysed using coupled gas chromatography-mass
spectrometry (Thermo Electron Trace GC-MS; 30 m DB5ms fused silica capillary column, 0.32 mm i.d., 0.25 lm
film thickness; carrier gas: helium). The usual GC temperature program was as follows: injection at 60°C; 2 min
isothermal; heating with a rate of 15°C min-1 to 150°C;
from 150 to 320°C at 4°C min-1, where it was held constant for 30 min. Identification of individual compounds
was achieved by comparison of mass spectra and GC
retention characteristics with published data. Quantification was based on adding compounds of known
concentration prior to analysis.
Compound-specific stable carbon isotope ratios were
measured using a HP5890 series II GC coupled via a
Thermo Electron GC-combustion-III-interface to a Finnigan MAT 252 mass spectrometer (GC conditions like
above). 13C/12C are reported in the delta notation relative
to the V-PDB standard (mean of at least three analytical
replicates). For polar compounds, values were corrected
for the addition of carbon during the derivatisation procedures. System precision was controlled using a mixture
of n-alkanes (C15–C29) with known carbon isotopic composition. Standard deviations were less than 0.4%.
7
Micritic layer,
11.10
rim around hole
A-1d
0.43
0.46
97.3
79.6
10.8
7.0
1.4
0.064
7.0
0.75
3.0318
NC
11.4
3.0019
NC
NC
11.71
12.49
3.0151
NC
9
Micritic layer
A-1c
7
Transition zone
A-1b
1.23
98.0
7.0
7.0
NC
NC
11.88
3.0151
NC
5
NC
Yellowish calcite
A-1a
5
IHMC/ICor d(104)HMC (Å) MgCO3 HMC ILMC/ICor d(104)LMC (Å) MgCO3 LMC ILMC/IHMC MgCO3 weighteda CaCO3 TOC C/N Remarks
(mol%)
(wt%) (wt%)
(mol%)
(mol%)
Gas chromatography and mass spectrometry
Sample ID Sample type
Table 1 XRD and geochemical bulk sediment analyses (CHN) on specific lithological facies
Int J Earth Sci (Geol Rundsch) (2009) 98:677–695
Interstitial water analyses
The sediment retrieved with the TV-grab was sampled for
interstitial water analysis using a push core driven into the
sediment when the TV-grab was still in its closed state.
The sediment in the pushcore was sampled in 2 cm
intervals and subsequently transferred into the cold room
(4°C) for further treatment. Pore fluids were retrieved by
squeezing using a pressure filtration system (0.2 lm cellulose acetate filters) at pressures up to five bars. The pore
fluid was separated immediately after extraction for
analyses. Pore water sub-samples were analysed for Clby Mohr’s titration (Grasshoff et al. 1999) directly
onboard with an analytical precision of 10 mM. Dissolved
SO2was determined by ion chromatography using a
4
Metrohm IC equipped with a conventional anion
exchange column and carbonate–bicarbonate solution as
solvent. The cations Mg2+ and Ca2+ were determined in
the shore-based laboratory at IFM-GEOMAR using an JY
170 ULTRACE inductively coupled plasma-atomic
emission spectrometer (ICP-AES) with an analytical
precision of 5%. Sub-samples for ICP-AES analyses were
acidified after pore water extraction to prevent mineral
precipitation due to degassing and depressurisation as
well as oxidation of reduced cations. Analyses for stable
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Int J Earth Sci (Geol Rundsch) (2009) 98:677–695
oxygen isotope composition of the pore water were performed at the GCA Laboratory in Sehnde (Germany)
using a PDZ Europa 2020 IRMS.
The d13C values of methane were determined as
described in Damm et al. (2007) by continuous-flow isotope ratio mass spectrometry (CF-IRMS) (Popp et al.
1995) at the Alfred Wegener Institute for Polar and
Marine Research (AWI) in Bremerhaven. The reproducibility derived from the duplicate measurement was
±0.5%. Isotopic ratios are reported relative to the Pee
Dee Belemnit (PDB) standard using conventional delta
notation
Results
Mineralogy
X-ray diffraction analyses and bulk inorganic carbon content measurements show that calcite is the dominant
mineral phase in all samples. It comprises 79.6 wt% in the
central part of the less cemented micritic layer (specimen
A-1), and much higher values were found in the coccolith
ooze (96.8 wt%), the yellowish calcite (98.0 wt%) and the
transition zone (97.3 wt%) (Table 1). Aragonite, dolomite
or Fe-bearing carbonates were not detected.
Samples containing low-magnesium calcite (LMC;
here defined as MgCO3 \ 8 mol%; [30 lm long needles;
Fig. 5b) as the dominating calcite phase are exclusively
found on the lower side of the carbonate precipitates and in
cavity fillings. High-magnesium calcite (HMC, MgCO3 [
8 mol%) is present in the upper parts of the samples (i.e.
the micritic layers and coccolith ooze) and is characterized
by small (\3 lm) micritic crystals (Fig. 5c). LMC additionally occurs in minor amounts in the coccolith ooze and
in the micritic layer in sub-sample A-1, which is due to the
presence of coccoliths (Figs. 5d, 6) that consist of LMC
(Trimonis 1974). Beside calcite, the diffractograms show
detrital minerals such as quartz, feldspars and clay minerals. As an exception, dark brownish coatings on the lower
side of the carbonate precipitates contain significant
amounts of framboidal pyrite and halite (‘‘pyritic coating’’
in Fig. 4).
Stable oxygen and carbon isotopes of carbonates
Stable oxygen and carbon isotope values scatter between
-40.8 and -31.5% V-PDB for d13C and -0.6 to +0.8 for
d18O (Table 2; Fig. 7a). In general, LMC samples have a
trend towards lower d18O values than HMC samples
(Fig. 7b), while for d13C measurements no obvious
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Fig. 6 X-ray diffractograms of different carbonate facies. The labels
(A-1a to B-2a) refer to the coding used in Table 1, numbers at the
peak tops display the computed MgCO3 content in mol% of the
dominating calcite (104) peak, calculated with the d(104) value after
Eq. 2. The inserts show an enlargement of the calcite (104) peak. a
LMC carbonates from the lower side of the carbonate structure. Note
the high NaCl content in B-1d, the pyritic coating. b HMC carbonates
from the middle and upper parts of the slabs. Note the LMC peak in
the coccolith-containing samples. The peaks at ca. 28.5° 2h are
artefacts, representing shadows of the prominent d(104) calcite peak
due to the use of an aged X-ray tube (Christoph Vogt, personal
communication)
correlations with the calcite composition exist. The d18O
vs. d13C cross plot (Fig. 7a) shows a tendency of apparently similar facies to cluster (e.g., the yellowish calcite
and the cemented coccolith ooze). Remarkable is the
correlation of d13C and d18O values measured for samples
from the micritic layers (Figs. 7a, 8): The most indurated
samples (from specimen B-1 and B-2 and the marginal
samples from the thick micritic layer in A-1) have the
highest d18O and lowest d13C values, and vice-versa for
the soft samples.
Int J Earth Sci (Geol Rundsch) (2009) 98:677–695
Table 2 Stable carbon and
oxygen isotope values of
carbonates, MgCO3, and
calculation of fluid temperatures
after Eq. 4
685
d18O (%)
MgCOb3
(mol%)
Piece
Facies
d13C (%)
Tcalc. (°C)a
d18Ow =
-1.77%
A-1
ML, rim (hole)
-35.89
0.34
11.5
9.50
6.47
A-1
ML, rim (hole)
-34.36
0.28
11.5
9.75
6.72
A-1
ML, rim (hole)
-33.86
0.18
11.5
10.18
7.13
A-1
ML, rim (hole)
-33.69
0.15
11.5
10.30
7.24
A-1
ML, rim (hole)
-33.68
0.14
11.5
10.34
7.29
A-1
ML, rim (hole)
-36.94
0.46
11.5
9.23
6.21
A-1
A-1
ML, rim (hole)
ML, rim (hole)
-38.47
-38.87
0.01
0.18
11.5
11.5
11.33
10.33
7.98
7.28
A-1
ML, upper rim
-38.38
0.44
11.5
9.31
6.29
A-1
ML, upper rim
-40.36
0.55
11.5
8.85
5.84
A-1
ML, centre
-40.23
0.47
10.8
9.19
6.17
A-1
ML, centre
-39.69
0.42
10.8
9.37
6.35
A-1
ML, centre
-38.85
0.50
10.8
9.03
6.02
A-1
ML, centre
-38.85
0.57
10.8
8.78
5.77
A-1
ML, centre
-38.83
0.57
10.8
8.76
5.76
A-1
CO
-32.05
0.45
9.5
8.74
5.74
A-1
CO
-35.61
0.53
9.5
8.41
5.42
A-1
YC
-37.94
-0.46
7.0
11.86
8.75c
A-2
ML, centre
-31.49
0.02
10.8
10.81
7.74
A-2
CO
-33.02
0.57
9.5
8.25
5.26
A-2
CO
-35.17
0.61
9.5
8.12
5.14
A-2
A-2
CO
YC
-33.18
-37.28
-0.22
-0.54
11.4
7.0
11.98
12.21
8.88
9.10c
A-2
YC
-37.67
-0.47
7.0
11.90
8.80c
A-2
TZ
-39.16
-0.17
7.0
10.67
7.61c
A-2
TZ
-33.53
-0.22
7.0
10.89
7.82c
B-1
ML
-40.63
0.63
11.3
8.46
5.46
B-1
ML
-40.30
0.63
11.3
8.47
5.48
B-1
ML
-40.41
0.74
11.3
8.02
5.04
B-1
ML
-39.65
0.42
11.3
9.32
6.30
B-1
YC
-32.78
-0.44
5.9
11.52
8.43c
B-1
YC
-38.42
-0.29
5.9
10.89
7.82c
Tcalc. (°C)
d18Ow =
-2.53%
B-1
Cavity fill
-38.23
-0.07
5.9
9.99
6.94
Value from Rank et al. (1999)
B-1
Cavity fill
-36.46
0.22
5.9
8.83
5.83
Weighted for LMC and HMC
content
B-1
CO
-35.49
0.32
11.4
9.76
6.72
B-1
Pyritic coating
-30.99
-0.38
5.9
11.26
8.19c
B-2
ML
-40.76
0.62
12.4
8.78
5.78
a
b
c
Value displayed in Fig. 7d
(diverging from Fig. 7c)
Bulk organic matter
Biomarker lipids
The organic carbon content (TOC) is relatively low in the
micritic layer and coccolith ooze (0.39–0.74 wt%) but
increases to 1.2 wt% in the yellowish calcite. The C/N
(C = TOC) ratio is between five and nine, a range typical
for marine organic matter (Holtvoeth et al. 2003).
Biomarker lipids characteristic for methanotrophic archaea
and sulfate-reducing d-proteobacteria involved in AOM
(Elvert et al. 1999; Thiel et al. 1999; Hinrichs et al. 2000;
Pancost et al. 2001; Thiel et al. 2001a; Pancost and Sinninghe Damsté 2003) are found in substantial amounts in
123
686
Int J Earth Sci (Geol Rundsch) (2009) 98:677–695
TZ
YC
-0.6
Cavity fill
13
a
MgCO3 (mol%)
δ18O (‰ V-PDB)
0
0.2
0.4
-42
13
-40
-38
-36
-34
δ13C (‰ V-PDB)
-32
8
LMC,
lower side
5
-0.6
-30
-0.4
-0.2
0
0.2
0.4
0.6
0.8
-32
-30
δ18O (‰ V-PDB)
13
c
12
calculated temperature (°C)
calculated temperature (°C)
9
6
0.8
11
10
9
8
d
12
11
10
9
8
7
7
-42
-40
-38
-36
-34
13
δ C (‰ V-PDB)
-32
-30
Fig. 7 a Crossplot of stable oxygen and carbon isotope measurements. Samples from the micritic layers (‘‘ML’’) show a separation
between values obtained from the soft centre part of the thick micritic
layer in piece A-1 (Fig. 4) and those from indurated, more cemented
micritic layers (ML in pieces B-1, B-2 and the rim of the thick ML in
A-1; Fig. 4). b Crossplot of d18Ocarbonate and MgCO3 content (mol%)
in the dominating calcite phase. Carbonates from the lower side of the
carbonate structures generally have lower MgCO3-contents than those
δ13C (‰ PDB)
10
7
0.6
-42
-40
-38
-36
-34
13
δ C (‰ V-PDB)
from the upper part. c Crossplot of d13C and calculated temperatures,
with a d18Owater of 1.77% (Rank et al. 1999) for all samples. d The
same as in c but calculated with a d18Owater of 2.53% (SMOW) for
the LMC samples from the bottom side of the carbonate slab
(Table 2). The error estimates are derived from the standard deviation
of the d18Owater measurements. Note that using this value the LMC
samples plot considerably closer to the present bottom water
temperature of 9°C (hatched line)
calcite in contrast to the coccolith ooze and the micritic
layer are evident.
Unit III
Sapropel
Coccolith ooze
-6
ML (soft)
HMC,
centre/
upper side
11
-0.2
ML (indurated)
b
12
-0.4
-8
CO
Pyritic coating
-4
y=-0.059x-1.79
r²=0.58
-2
Yellowish calcite
0
2
-44
-40
-36
-32
-28
-24
-20
-16
-12
-8
-4
0
δ13C (‰ V-PDB)
Fig. 8 d18O and d13C measurements from the micritic layers (open
diamonds indurated, closed diamonds soft) showing a linear correlation. The continuation of the regression line plots between d18O
measurements of diagenetically unaltered bulk sediment (i.e., coccolith ooze stars, sapropel triangles, and lacustrine Unit 3 circles),
indicating the admixtures of detrital and biogenic carbonates from
different strata. The d18O data obtained on bulk sediment are from
Deuser (1972) and Major et al. (2002)
all fractions prepared from the major lithological facies of
the carbonate precipitate (Table 3). However, significant
differences in the distribution patterns of the yellowish
123
Non-isoprenoidal dialkylglycerol diether lipids (DAGE)
with an inferred sn-1,2 stereochemistry and 32–37 carbon
atoms are the prevailing compound class of the extractable
alcohols in the yellowish calcite (Table 3), and a DAGE
with two 12-methyltetradecane (anteiso, ai-C15) moieties is
the major component of the fraction. These compounds are
commonly attributed to sulfate-reducing d-proteobacteria
(Pancost and Sinninghe Damsté 2003). Other lipids
attributed to bacteria in the yellowish calcite are shortchain n-alcohols (C14–C20) and, only in trace amounts,
non-isoprenoidal mono alkyl glycerol monoether lipids
(MAGE). The MAGE contain C14–C18 alkyl chains which
are ether-bound putatively at sn-1, and are dominated by
the ai-C15 compound. Archaeol (2,3-di-O-phytanyl-sn-glycerol)
Int J Earth Sci (Geol Rundsch) (2009) 98:677–695
687
Table 3 Concentrations (lg g-1 carbonaceous material dry weight)
and d13C values of selected compounds from diverse source
organisms as well as concentrations of compound classes (sum of
Micritic layer
-1
Conc. (lg g )
single compounds) retrieved upon dissolution of the three main
lithological facies constructing carbonate sample A-2
Coccolith ooze
13
d C (%)
-1
Conc. (lg g )
Yellowish calcite
13
d C (%)
Conc. (lg g-1)
d13C (%)
Prokaryotic lipids (putatively AOM-associated)
Crocetane
0.005
NA
0.007
NA
0.133
-97.5
PMID0
0.053
NA
0.122
NA
0.234
-107.8
i-C15:0-SCFA
0.050
NA
0.669
NA
2.980
-99.3
ai-C15:0-SCFA
0.151
NA
1.238
NA
10.327
-92.5
ai-C15:0-MAGE
ND
ND
0.023
NA
0.073
NA
ai-C15:0/ai-C15:0-DAGEb
0.628
- 98.3
2.680
-89.7
6.557
-94.3
Archaeol
0.467
- 88.7
2.614
-68.0
4.314
-94.3
sn-2-OH-archaeol
ND
ND
0.790
NA
1.424
-99.0
R PMID1-4a
0.100
0.338
0.822
R MAGE
0.037
0.345
0.511
R DAGE
2.221
8.384
16.043
Ratios
SCFA:MAGE:DAGE
43:1:60
20:1:24
56:1:32
OH-Ar-indexc
PMI D0/R PMED1-4a
NC
0.53
0.24
0.36
0.35
0.29
Eukaryotic lipids
Cholesterol
0.098
NA
0.178
NA
0.057
NA
-31.0
1.297
NA
ND
ND
Dinosterol
0.267
R diols (C28, C29, C31)
0.461
3.151
0.032
R keto-ols (C30, C32)
0.304
1.451
0.200
R n-LCFAs (C22–C30)
R n-LCHCs (C25–C31)
1.130
0.212
1.082
0.221
0.334
0.043
R n-LCOHs (C23–C30)
0.429
0.531
0.116
Lipids from potentially diverse sources
Ratios
n-C16:0/n-C16:1 SCFA
4.45
2.55
0.76
n-C18:0/n-C18:1 SCFA
1.39
0.51
0.26
NA not analysed, ND not detected, NC not calculated, PMID0 = 2,6,10,15,19-pentamethylicosane, PMID1-4 = pentamethylicosenes with up to
four double bonds, cholesterol = cholest-5-en-3b-ol, dinosterol = 4a,23,24-trimethyl-5a-cholest-22-en-3b-ol, SC = short-chain, LC long-chain,
MAGE monoalkyl glycerol monoethers, DAGE dialkyl glycerol diethers, FA fatty acid, OH alcohol, HC hydrocarbon
a
Partly co-elutions with n-C23:1 and n-C23:0 (Thiel et al. 2001b)
b
Co-elution with small amounts of DAGE emcompassing i- and/or ai- branched C15:0 moities
c
The OH-Ar index gives the ratio of sn-2-hydroxyarchaeol to archaeol [(archaeol + sn-2-hydroxyarchaeol)/archaeol - 1] (Blumenberg et al.
2004)
followed by sn-2-hydroxyarchaeol (2-O-3-hydroxyphytanyl-3-O-phytanyl-sn-glycerol) are the major archaeal lipids
in the alcohol fraction. Small amounts of unidentified
unsaturated archaeol derivatives as well as a C40 macrocylic diether with one cyclopentyl ring, previously reported
by Stadnitskaia et al. (2003), are also present. The hydrocarbon fraction of the yellowish calcite is characterized
by high amounts of acyclic isoprenoidal compounds.
2,6,10,15,19-pentamethylicosane (C25, PMID0) and its
mono- to tetraunsaturated counterparts (pentamethylicosenes,
PMID1-4) are by far the dominating structure equivalents
of the fraction. The C20 homologue crocetane (2,6,11,15tetramethylhexadecane) was found in moderate amounts,
and crocetenes were not detected. In the fatty acid fraction
of the yellowish calcite, saturated and monounsaturated
short-chain fatty acids (SCFA) with 12–20 carbon atoms
123
Coccolith ooze
All bacterial and archaeal biomarkers found in the yellowish calcite are also present in the coccolith ooze,
although in varying abundances. Notably, saturated and
unsaturated homologues of PMI predominate the hydrocarbons of the coccolith ooze, whereas crocetane was only
found in trace amounts. However, compounds commonly
attributed to planktonic organisms, like a suite of C27 to
C29 sterols including cholesterol (cholest-5-en-3b-ol) and
dinosterol
(4a,23,24-trimethyl-5a-cholest-22-en-3b-ol),
long-chain 1,15-diols (C28, C29, and C31) and the structural
equivalent 1,15-keto-ols (C30 and C32), are present in much
higher abundances relative to those detected in the yellowish calcite (Table 3). Further saturated long-chain nalkyl compounds, like C22–C30 fatty acids, C22–C30 alcohols, and C25–C31 hydrocarbons, indicative of inputs from
higher plants (Eglinton et al. 1962; Simoneit 1977), were
found in high amounts in the respective lipid fractions
prepared from the coccolith ooze.
Micritic layer
In the micritic layer, the distribution patterns of all lipid
fractions are similar to those of the coccolith ooze. The
most obvious contrasts in comparison to the coccolith ooze
are lower abundances of lipids attributed to AOM performing microorganisms and significant depletions in
unsaturated compounds relative to their saturated homologues, as exemplified for PMI derivatives and n-C16 and
n-C18 SCFA (Table 3).
C l - (m M )
S O 4 2 - (m M )
0 2 4 6 8 10 320
depth (cmbsf)
prevail. Major components are odd-chain terminal methyl
branched compounds with 15 (i- and ai-C15:0-SCFA) and
17 carbon atoms.
Remarkably, compounds attributed to AOM performing
prokaryotes and present in amounts sufficient for stable
carbon isotope ratio measurements in the three fractions
exhibit extreme depletions in 13C with d13C -68.0 to 107.8% (Table 3).
Int J Earth Sci (Geol Rundsch) (2009) 98:677–695
328
336
0
0
5
5
10
10
15
15
20
20
3.2
3.6
M g /C a
4
-2.8
-2.6
depth (cmbsf)
688
-2.4
δ 1 8 O (‰ SMOW)
Fig. 9 SO2concentration, Mg/Ca ratio, Cl- content and d18O of
4
interstitial water from station BS346GR. Note that a certain scatter
might be produced due to sampling potentially disturbed sediment
down to 3.23 at 21 cm depth, while chloride concentration
is dropping likewise from 335.8 to 323.1 mM. The d18O
measurements scatter between -2.72 and -2.43%. The
values for the measured parameters of the uppermost sediments differ from the present-day values for Black Sea
bottom water [sulphate-concentration: 18 mM (Murray
et al. 1991); chloride-concentration: 350 mM (Aloisi et al.
2004); Mg/Ca ratio: 4.77 mol/mol (Manheim and Chan
1974); d18O: -1.77% (Rank et al. 1999)], likely due to
loss of the top few centimeters of sediment during core
recovery. Since the samples were taken with a pushcore
from a freshly recovered and unopened TV grab the results
might be subject to some disturbances, but nevertheless
provide useful information about the general trends in the
interstitial water, especially when compared to the Black
Sea bottom water.
The methane dissolved in interstitial water at shallow
sediment depth shows a strong depletion in 13C
(d13Cmethane = -79.8 and -80.8%, n = 2) indicating its
microbial origin.
Discussion
Interstitial water
Characterisation of carbonate formation conditions
from mineralogical and stable isotope data
Reported are interstitial water profiles from station
BS346GR covering the uppermost 21 cm of the retrieved
sediment column (Fig. 9). Sulphate concentrations
decrease from 10 mM at the surface approaching 0 mM at
21 cm. The molar Mg/Ca ratios also decline from 3.87 at
1 cm core depth (the sample from 0 cm might be an artifact due to disturbances during the sampling procedure)
In recent years, several studies (e.g., Mazzini et al. 2004;
Michaelis et al. 2002; Peckmann et al. 2001; Pape et al.
2005; Reitner et al. 2005a; Stadnitskaia et al. 2005)
described the occurrence of authigenic carbonates in association with methane (micro-) seepage in shallow (oxic) and
deep (anoxic) waters of the Black Sea. Typical for these
AOM-related carbonates are strongly negative d13C values
123
Int J Earth Sci (Geol Rundsch) (2009) 98:677–695
689
that also characterize the samples from the DM (Fig. 7a).
An average d13C value of -80.3% measured for methane
extracted from interstitial waters in sediments retrieved by
the TV-grab, indicates a biogenic origin of the methane at
the DM site (Whiticar 1999).
One of the remarkable characteristics of the carbonate
samples from the DM is the spatial discrimination of HMC
against LMC within the carbonate slabs. This slightly
resembles the aragonite/HMC zonation described from
carbonate towers found in the NW Black Sea (Reitner et al.
2005b), which was interpreted to be ultimately governed by
the sulphate gradient of the ambient water. If sulphateconcentration would also be the driving factor that determines whether LMC or HMC is formed at the DM site,
LMC should be expected on the upper part of the slabs:
higher sulphate concentrations suppresses the Mg-incorporation into calcite (Burton 1993) and sulphate
concentration declines in interstitial waters with depth
(Fig. 9). Instead, LMC is found on the lower side of the
carbonate structures. However, the observation that the
dominant calcite in the overlying coccolith ooze has lower
MgCO3 contents (10.3 mol% MgCO3) compared to the
deeper positioned micritic layers (11.3–12.4 mol%
MgCO3) (Table 1; Fig. 6) could be tentatively ascribed to
the sulphate profile of the interstitial water.
Other factors that are important for the Mg incorporation
into calcite are the Mg/Ca ratio and the temperature of the
fluid, where both, higher Mg/Ca and temperature, lead to
higher Mg contents in the calcite (e.g., Hartley and Mucci
1996; Morse et al. 2007).
Increased fluid temperatures would cause an increased
Mg uptake into calcite (see, e.g., Hartley and Mucci 1996).
Calcite formation temperatures can be calculated after
Friedman and O’Neil (1977):
d18 Ocalcite ðV-PDBÞ ¼ e
2:78ð106 T 2 Þ2:89þ0:06Mg%
1;000
1; 000 þ d18 Owater 1; 000
ð4Þ
where d18Ocalcite is the measured stable oxygen isotopic
composition of the calcite, d18Owater is the isotopic composition of the ambient water or fluid (in % V-PDB), T the
fluid temperature in Kelvin and Mg% the MgCO3 content
of the calcite in mol% after Eq. 2. If HMC and LMC
phases occurred simultaneously (Table 1), a weighted
mean was taken for MgCO3% based on the ratio of the
intensities of the respective calcite (104) peaks.
The results (Table 2; Fig. 7c) show that if the d18Owater
for Black Sea bottom water [-1.77% SMOW (Rank et al.
1999)] is used, the temperatures calculated after Eq. 4
scatter generally between 8.0 and 12.2°C. Especially the
LMC-samples from the lower part of the precipitates
(yellowish calcite, transition zone, pyritic coatings) have
calculated formation temperatures significantly higher than
the present bottom water temperature of about 9°C. Since
the LMC-samples are located at the base of the carbonate
precipitates, they are more likely to precipitate in equilibrium with interstitial water. Reasonable temperature
estimates for the respective LMC-samples were calculated
with a d18Owater of -2.53% (SMOW) (Fig. 7d; Table 2),
which is the average of the d18Owater values measured on
interstitial water at the sampling site BS346GR. Although
the exact depth of the carbonate slabs is not well-constrained and the interstitial water profile scatters
considerably, these calculations indicate that the LMC and
HMC samples precipitated under different environmental
conditions also affecting the elemental composition of the
interstitial water as discussed below.
Slightly higher calculated formation temperatures for
the carbonate of the micritic layers are due to the admixture
of detrital calcium carbonate, particularly coccoliths, that
carry a sea surface temperature signal (note also the slight
LMC-peak in the XRD data of the micritic layer subsample A-1e in Fig. 6b). In conclusion, there is no indication for any distinguished temperature gradient at the
DM site, which would explain at least partly the LMC
versus HMC partitioning.
The observed trend towards lower Mg/Ca ratios of the
interstitial water in greater depths (Fig. 9) is in line with
LMC precipitation at the lower side of the carbonate slabs
(Morse et al. 2007), but the observed Mg/Ca gradient is
quite small and might not be sufficient to explain the distinctive LMC–HMC discrimination in our samples.
However, seeps are often subject to considerable changes in the rate of fluid flow, which leads to substantial
variations in shape of the interstitial water profile. Temporal changes related to sedimentation and diagenetic
processes have also to be taken into account: ongoing
cementation will clog the initially available pore space
within the hemipelagic sediment, thereby restricting the
methane flux and AOM rates and, hence, reducing the
HMC precipitation. If the mineralogy and geochemical
properties of the micritic HMC are characteristic of an
earlier stage of the carbonate slab evolution, the HMC
should have been precipitated under the influence of
interstitial water with a composition similar to Black Sea
bottom water, since the sediment has been closer to the sea
floor than at present (continuous hemipelagic sedimentation presumed). This assumption is supported by the
previous temperature calculations, which indicate that the
HMC cementing the coccolith ooze and micritic layers has
apparently been precipitated in equilibrium with the Black
Sea bottom water. It follows that recent active calcite
growth is mostly confined to the yellowish LMC rims on
the lower side of the precipitates where the microbial mats
are thriving and no direct contact to bottom water is provided. It is important to note that both, Mg/Ca ratio and
123
690
sulphate concentrations, increase towards the seafloor but
both parameters are antagonistic in their effect on the Mgcontent of the precipitated calcite. It seems that in this
particular case the influence of the Mg/Ca ratio might have
outweighed the influence of the sulphate concentration.
The temperature calculations based on the measured
d18Ocarbonate values also indicate that at the sampled site
gas hydrates were not present in significant amounts,
although the DM site lies within the gas hydrate stability
zone based on the present bottom water temperature and
chlorinity of the deep Black Sea [about 9°C, chlorinity = 350 mM (Aloisi et al. 2004)] (Fig. 10). Hydrates
incorporate preferably the heavy isotope 18O, thus active
hydrate formation will lead to 18O-depletion of the
remaining interstitial water, while its decomposition would
release a fluid that is enriched in 18O by up to +2.9%
(SMOW) (Hesse 2003). The chlorinity profile (Fig. 9) also
shows no excursions that might imply gas hydrate dissociation (producing negative anomalies) or active formation
(positive anomalies) (Haeckel et al. 2004). The trend
towards lower chlorinity, Mg/Ca, and d18O values of the
interstitial water with greater depth argues in favour of an
advection of low-salinity fluids from the lacustrine sediments. Thus, no direct or indirect evidence for the presence
of gas hydrates in the vicinity of the carbonate precipitates
exists. Methane hydrate might have formed in greater depth
beneath the retrieved carbonates and thus left no imprint in
our data from relatively shallow depths.
Fig. 10 Methane hydrate stability diagram for the Dolgovskoy
Mound position (after Bohrmann et al. 2003; Sloan 1998). Black
Sea water chlorinity is 350 mM, bottom water temperature 9°C, the
temperature gradient below the sea floor is assumed to be 29°C/km.
GHSZ gas hydrate stability zone, BGHSZ base of gas hydrate stability
zone
123
Int J Earth Sci (Geol Rundsch) (2009) 98:677–695
Relation of biomarker lipids to the different lithological
facies
Lipid biomarker analyses gave valuable information on the
prevailing AOM communities being involved in the precipitation of authigenic calcite at the DM. Based on the
prevalence of archaeol over sn-2-hydroxyarchaeol (OH-AR
index = 0.24–0.35), and the ‘‘DAGE over MAGE and fatty
acid predominance’’ (Table 3), an ANME-1 dominated
AOM consortium (Blumenberg et al. 2004) is inferred for
all sub-samples. However, deduced from substantial
amounts of crocetane and sn-2-hydroxyarchaeol present in
the yellowish calcite, contributions by ANME-2 archaea
can not be ruled out.
In two separate studies collective lipid biomarker and
molecular biological approaches were used to elucidate the
main archaeal community involved in the precipitation of
methane-related carbonate precipitates at deep-sea mud
volcano (MV) settings. Aloisi et al. (2002) reported
crocetane/PMID0 ratios similar to those observed for our
yellowish calcite and the occurrence of ANME-1 specific
16S rDNA sequences in a carbonate crust (MN16BT2)
from the Napoli MV (Eastern Mediterranean), while
sequences belonging to the ANME-2 group could not be
detected. Likewise, the exclusive presence of ANME-1
related 16S rDNA genes in two carbonate crusts sampled at
the NIOZ MV in the Sorokin Trough (Black Sea) was
described by Stadnitskaia et al. (2005). For both crusts
from the Sorokin Trough, a co-occurrence of archaeol and
hydroxyarchaeol was observed. These observations suggest
that the production and incorporation of ANME-1-derived
crocetane and hydroxyarchaeol in our samples can also not
be excluded.
The apparent absence of biomarkers attributed to
organisms being not involved in the AOM, together with a
strong depletion in 13C observed for all compounds analysed for the yellowish calcite suggests that the material
was exclusively constructed by methanotrophic archaea
and sulphate-reducing bacteria growing on inner surfaces
of cavities. Consequently, these cavities should have been
filled, both, with methane and sulphate-bearing fluids in
amounts adequate to maintain ANME-1 communities.
Voids and cracks within the carbonates might serve as
main pathways for the downward diffusion/advection of
sulphate to the cavities since the porosity of the sediments
is reduced by carbonate cementation.
In contrast, the admixture of non-autochthonous components in the micritic layer and coccolith ooze located
above the yellowish calcite is evidenced by a high proportion of lipid biomarkers attributed to planktonic and
terrestrial organisms, like long-chain n-alkyl compounds
(Eglinton et al. 1962; Simoneit 1977), sterols (Volkman
1986), as well as long mid-chain diols and keto-ols
Int J Earth Sci (Geol Rundsch) (2009) 98:677–695
(Versteegh et al. 1997). Based on the relative abundances
of individual lipids, the coccolith ooze sample appears to
represent to some extent a mixture of compounds also
found in the yellowish calcite and in the micritic layer.
These results correspond to the mineralogical and geochemical data which show that the yellowish calcite is a
pure authigenic carbonate while both, the micritic layer and
the coccolith ooze, contain some amounts of detrital
components, like coccoliths (see also Figs. 5d, 6b). In
particular for the yellowish calcite high abundances of
PMID1-4 and monounsaturated SCFA relative to their
saturated counterparts were observed (Table 3). Regardless
of different source organisms (bacteria and eukaryotes)
potentially contributing SCFAs to the three differents
lithological facies, these findings suggest that comparatively unaltered organic material was incorporated and,
thus, protected against biotic and abiotic alterations.
The lipid patterns as well as the carbonate mineralogy
found in our samples strongly resemble those observed for
AOM-mediating mats and associated lateral extensive
carbonate precipitates from other parts of the Black Sea
like the Ukrainian shelf, the central Black Sea and the
Sorokin Trough (Mazzini et al. 2004; Pape et al. 2005;
Stadnitskaia et al. 2005), and other AOM environments.
However, our results differ from the findings made for the
tall carbonate towers from the northwestern Black Sea
slope, where distinctly different carbonate phases co-occur
with different AOM consortia (Reitner et al. 2005b). In our
samples this is not the case, both, the HMC and LMC parts
of the carbonates inherit the lipids of similarly structured
AOM performing consortia, namely ANME-1 dominated
communities. This is somewhat surprising since the
microenvironments in which the organisms thrive are quite
different: on one hand they are growing as thick microbial
mats into the fluid-filled cavities where a higher methane
availability can be assumed, on the other hand they are
forming thin biofilms along the laminations and microvoids in the coccolith ooze and micritic layer (Mazzini
et al. 2004) where a diffusive fluid regime is prevailing and
thus a lower methane partial pressure is expected. Our
results from the DM indicate furthermore that the initial
(micro)structure and mineralogy of the sediment does not
seem to be a crucial factor for the composition of AOMperforming consortia inhabiting the subsurface.
Origin of the micritic layers
The micritic layer(s) might have played a crucial role in
shaping the carbonates, through acting as a barrier for the
uprising fluid. It is assumed that the subsequent build-up of
overpressure beneath these layers created the large cavities
forming on the lower side of the carbonate structures while
691
deforming the overlying sediments when they were still not
fully calcified. Getting insight into the origin of the micritic
layers might therefore be important for understanding the
formation of the carbonate precipitates.
Similar homogeneous and structureless clayey layers
with limited lateral extension have been previously described as ‘‘Degens layers’’ (Kempe et al. 2001). The lack of
lateral continuity argues against turbidite currents as processes creating the layers. Kempe et al. (2001) attributed
them instead to the expulsion of overpressurized interstitial
water and mud that penetrated the less permeable sapropel
and coccolith ooze. Ascending fluids could produce, or at
least enhance, occasional spouts of liquidized mud, which is
then deposited on a spatially restricted area and subsequently cemented by methane-derived HMC. If this model
is correct, this would also require several periods of
enhanced fluid and/or gas emission flow following relatively shortly after each other, divided by just a few
coccolith ooze laminae, which represent approximately a
year each (Hay et al. 1991) (c.f. the succession of three
micritic layers in B-1, Fig. 4). While the model of Kempe
et al. (2001) implies that the material making up the micritic
layers consists to a huge degree of sediment derived from
the lacustrine Unit 3, our data suggests that the primary
detrital material includes also large proportions of reworked
coccolith ooze and sapropel. As shown in Fig. 8 stable
carbon and oxygen isotopes of samples from the micritic
layer follow a linear trend that continues towards values
between those measured in previous studies for samples of
the unlithified coccolith ooze, sapropel and Unit 3 (Deuser
1972; Major et al. 2002). This indicates that the bulk
d13Ccarbonate and d18Ocarbonate signal from the micritic layers
is a mixture between AOM-induced HMC and the detrital
and biogenic carbonates in the hemipelagic sediment, which
apparently includes material from the bordering coccolith
ooze but also from the stratigraphic older sapropel and/or
lacustrine mud. The siliciclastic components in the diffractograms give more complex information: here the noncarbonaceous components of the thin micritic layers of
carbonate piece B are very similar in relative abundance
compared with those from the underlying coccolith ooze in
piece B. This is different for the more massive micritic layer
in piece A which includes clay minerals, quartz and plagioclase which are barely detectable in the coccolith ooze
sample from the same piece A (Fig. 6). Thus, it seems as if
the material from smaller layers is derived from directly
underlying strata, while the thicker layers represent ‘‘microeruptions’’ that reactivated material from a greater range of
depths (sapropel and upper Unit 3). As observed in the thick
micritic layer of piece A-1 (Fig. 4) the carbonate cementation following the deposition of the layers is directed from
the outer parts (rim) of the micritic layer, where fluid flow is
most intense, towards the centre.
123
692
Int J Earth Sci (Geol Rundsch) (2009) 98:677–695
b Fig. 11 Sequence of events leading to the formation of carbonate
precipitates at the Dolgovskoy Mound (modified after Mazzini et al.
2007). 1 Advection of methane-saturated fluids causes the expulsion
of a layer of liquidized, homogenous clay that is reworked from the
upper sediment column. 2 Lateral migration of the fluids fuels AOM
and calcite precipitation in the upper part of the sedimentary column.
The calcite-cemented clay layer represents now a ‘‘micritic layer’’. 3
Reduced permeability and additional sealing through the covering
clay layer impedes the ascent of the fluids and leads to the formation
of fluid-filled cavities and subsequent deformation of the sediments.
Microbial mats develop at the surfaces of the cavities. 4 Overpressurized fluids break through the sediment cover, creating (sub)
vertical conduits. The yellowish calcite forms where thick microbial
mats are situated (at the voids and cavities). The deformation stops
when the sediment is fully calcified. Since the vertical migration of
fluids is largely blocked now, lateral fluid migration dominates and
causes the lateral expansion of the carbonate precipitates. The grey
shading indicates the calcified parts of the sediment, that constitute
the sampled carbonates. The scale is an approximation and meant for
spatial orientation
(2)
(3)
(4)
Chronological succession of events leading to the
formation of carbonate precipitates at the DM
Based on the conceptual model proposed by Mazzini et al.
(2007) we suggest that the different lithological facies
forming the carbonate structures at the DM have been
created by a complex interaction of sedimentological,
microbiological and geochemical factors (Fig. 11):
(1)
During a period of enhanced fluid flow, liquidized
homogenous clay is spilled out on a spatially
restricted area on the sea floor. This might happen
123
several times, creating a suite of clayey layers
separated by laminae of coccolith ooze (Unit 1).
Lateral fluid migration along the coccolith ooze
laminae fuels AOM and therefore authigenic HMC
precipitation (as described in Mazzini et al. 2004),
affecting also the clay layers (now called ‘‘micritic
layer’’). Since the calcite precipitation occurs near the
sea floor, the stable oxygen isotopic signature of the
HMC is mainly derived from the Black Sea bottom
water.
Increasing fluid/gas pressure deforms the still plastic
sediments and creates fluid/gas-filled cavities. On the
surfaces of the cavities thick microbial mats start to
thrive and the yellowish LMC rims precipitate, with
d18Ocarbonate values in equilibrium with the interstitial
water.
When the slabs are still not fully cemented, fluids
might be released through (sub)-vertical voids. The
predominance of ANME-1 dominated AOM-associations, which appear to be better adapted to low
methane partial pressure (Nauhaus et al. 2005),
suggests that the apparently high fluid pressure
necessary to create the big cavities has prevailed
only temporarily. The clogging of the pore space in
the upper part of the slabs reduces methane availability and therefore AOM rates.
Conclusions
Authigenic carbonates at the DM from the deep northeastern Black Sea have been formed by the interaction of
fluid seepage and the activity of microbial consortia
performing the AOM, which causes extensive calcite
Int J Earth Sci (Geol Rundsch) (2009) 98:677–695
precipitation. The AOM-consortia, dominated by the
ANME-1 group, form voluminous microbial mats within
the fluid-filled cavities and on the lower side of the carbonate precipitates. Significant discrepancies between the
precipitated authigenic calcites with respect to their magnesium content and stable oxygen isotopic composition
have been found, mainly discerning between micritic HMC
cementing the hemipelagic sediment and LMC forming
compact yellowish calcite rims in direct contact to the
microbial mats. The different geochemical and mineralogical characteristics of the carbonates are primarily a
result of different interstitial water chemistries prevailing
in the different sediment depths where the respective carbonate phases precipitated. The lipid biomarker
composition determined for the three different lithological
facies suggests that the structure of the AOM-performing
communities is very similar within the whole carbonate
structure, independent of lithology and microenvironment
(especially methane availability).
Acknowledgments We thank captain and crew of the R/V Professor Logachev for their excellent support while handling TV-grab and
gravity coring during the TTR-15 cruise. We highly appreciate the
help provided by many people, namely Christoph Vogt for doing the
XRD measurements, Monika Segl for stable isotope measurements,
Hella Buschhoff for CHN analyses and Hartmut Mai for running the
SEM/EDX, Markus Elvert and Daniel Birgel for support during lipid
biomarker analyses (all Univ. Bremen), Bettina Domeyer, Regina
Surberg (IFM-GEOMAR) for retrieving and analysing pore water
samples, Ellen Damm (AWI) for providing d13C measurements on
methane and Katja U. Heeschen (NOC, Southampton) for on-board
gas sampling. d18O measurements on interstitial water were performed by Manfred Schmitt (GCA, Sehnde), who is gratefully
acknowledged. Janis Thal and Svenja Papenmeier (Univ. Bremen)
provided very valuable support in the early stages of this research
project. This paper has benefited from constructive reviews by Barbara Teichert and Giovanni Aloisi. This is contribution GEOTECH286 of the R&D-programme GEOTECHNOLOGIEN funded by the
German Ministry of Education and Research (BMBF) and the German Research Foundation (DFG), collaborative project METRO
(grant 03G0604A). RCOM 0526.
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