GeolCarp_Vol66_No5_409

www.geologicacarpathica.com

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

, OCTOBER 2015, 66, 5, 409—426 doi: 10.1515/geoca-2015-0034

Depositional environment, organic matter characterization

and hydrocarbon potential of Middle Miocene sediments

from northeastern Bulgaria (Varna-Balchik Depression)

ALEXANDER ZDRAVKOV

, ACHIM BECHTEL

, STJEPAN ĆORIĆ

and

REINHARD F. SACHSENHOFER

Department of Economic Geology, University of Mining and Geology “St. Ivan Rilski”, 1700 Sofia, Bulgaria; alex_zdravkov@mgu.bg

Department Angewandte Geowissenschaften und Geophysik, Montanuniversität Leoben, Peter-Tunner-Str. 5, A-8700 Leoben, Austria;

achim.bechtel@unileoben.ac.at; reinhard.sachsenhofer@unileoben.ac.at

Geological Survey of Austria, Neulinggasse 38, A-1030 Vienna, Austria;

stjepan.coric@geologie.ac.at

(Manuscript received January 9, 2015; accepted in revised form June 23, 2015)

Abstract: The depositional environments and hydrocarbon potential of the siliciclastic, clayey and carbonate sedi-
ments from the Middle Miocene succession in the Varna-Balchik Depression, located in the south-eastern parts of
the Moesian Platform, were studied using core and outcrop samples. Based on the lithology and resistivity log the
succession is subdivided from base to top into five units. Siliciclastic sedimentation prevailed in the lower parts of
units I and II, whereas their upper parts are dominated by carbonate rocks. Unit III is represented by laminated clays
and biodetritic limestone. Units IV and V are represented by aragonitic sediments and biomicritic limestones, corre-
lated with the Upper Miocene Topola and Karvuna Formations, respectively. Biogenic silica in the form of diatom
frustules and sponge spicules correlates subunit IIa and unit III to the lower and upper parts of the Middle Miocene
Euxinograd Formation. Both (sub)units contain organic carbon contents in the order of 1 to 2 wt. % (median: 0.8 for
subunit IIa; 1.3 for unit III), locally up to 4 wt. %. Based on Hydrogen Index values (HI) and alkane distribution
pattern, the kerogen is mainly type II in subunit IIa (average HI = 324 mg HC/g TOC) and type III in unit III (average
HI ~ 200 mg HC/g TOC). TOC and Rock Eval data show that subunit IIa holds a fair (to good) hydrocarbon genera-
tive potential for oil, whereas the upper 5 m of unit III holds a good (to fair) potential with the possibility to generate
gas and minor oil. The rocks of both units are immature in the study area. Generally low sulphur contents are prob-
ably due to deposition in environments with reduced salinity. Normal marine conditions are suggested for unit III.
Biomarker composition is typical for mixed marine and terrestrial organic matter and suggests deposition in dysoxic
to anoxic environments.

Key words: organic geochemistry, hydrocarbon potential, Euxinograd Formation, NE Bulgaria, Eastern Paratethys.

Introduction

The western Black Sea has long been recognized as a region
with high petroleum potential. The Tyulenovo Field (Fig. 1b)
was detected in the Bulgarian sector of the Black Sea in the
1950s. It contains significant oil ( ~ 30 Mbs) and gas reserves
( ~ 30 Bcf;  Georgiev  2012).  Additional  potentially  oil-  and
gas-bearing  structures  have  been  identified  on  and  offshore
Bulgaria in the 1990s. One of them, the Galata deposit, off-
shore  cape  Galata  near  Varna,  already  produced  dry  gas  of
about 55 Bcf (Georgiev 2012).

The question of active source rocks is an important issue

for understanding petroleum systems. Georgiev (2012) con-
sidered the Middle to Late Jurassic shales from the Etropole
and  Provadia  Formations,  drilled  in  the  Galata  area,  as  the
most likely source rocks for the formation of Tyulenovo oil.
The same author, however, reports the presence of oleanane,
which  is  an  indicator  for  post-Jurassic  source  rocks,  within
this  oil.  Sachsenhofer  et  al.  (2009)  studied  Oligocene  clays
of the Ruslar Formation and determined a good potential for
hydrocarbon  generation.  Apart  from  Oligocene  rocks,
Eocene  successions,  which  are  of  considerable  thickness  in

the deeper parts of the western Black Sea may hold a source
potential (Georgiev 2012).

No studies have been conducted so far on the Miocene

succession,  although  there  are  some  notes  on  the  presence
of organic matter-rich lenses within the diatomaceous clays
of  the  Middle  Miocene  Euxinograd  Formation  (Koleva-
Rekalova 1997). Thus, the limited knowledge on the amount
and  origin  of  the  organic  matter  within  the  Miocene  rocks
leaves a gap in our understanding of the petroleum system in
NE Bulgaria and its off-shore part.

Therefore, the main aim of the present paper is to study li-

thology, depositional environment, and hydrocarbon potential
of the Miocene Euxinograd Formation and to explore vertical
and  partly  lateral  facies  variations.  To  reach  this  goal,  core
material from Miocene rocks from borehole C-180a (Fig. 1b)
was  studied.  Additional  samples  from  the  upper  part  of  the
Euxinograd  Formation  were  taken  from  an  outcrop  near
Albena  resort.  The  present  investigation  is  based  on  bulk
geochemical  parameters  (inorganic  and  organic  carbon,  sul-
phur,  hydrogen  index),  biomarker,  and  stable  carbon  isotope
composition of the organic matter. Since the lower part of the
Miocene succession, especially in the Balchik area, is not well

410

ZDRAVKOV, BECHTEL, ĆORIĆ and SACHSENHOFER

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

studied due to limited core material, we also investigate rocks
underlying the Euxinograd Formation.

Geological setting

Miocene rocks are present in the southeastern part of the

Moesian  Platform  in  Bulgaria  and  Romania  (Popov  et  al.
1986;  Popov  &  Kojumdgieva  1987;  Fig. 1)  and  have  been
studied intensively, because of their rich macro- and micro-
fauna,  as  well  as  diatom  flora  (Kojumdgieva  1965;
Kojumdgieva & Dikova 1978; Kojumdgieva & Popov 1987;
Darakchieva 1989; Temniskova-Topalova 1990).

The Miocene succession, about 300 m thick, has been depos-

ited in the relatively small, south-east widening Varna-Balchik
Depression (Fig. 1c) of the Eastern Paratethys (Kojumdgieva
& Popov 1981). The sediments overlie denudated Late Oli-
gocene clays (Ruslar Fm) and represent the time span from
the early Middle Miocene (Tarkhanian) to the middle Late

Miocene (Khersonian) (Popov & Kojumdgieva 1987). Based
on differences in sedimentary facies, Kojumdgieva & Popov
(1981), Popov et al. (1986) and Popov & Kojumdgieva (1987)
subdivided the Varna-Balchik Depression into the Varna
part, dominated by shallow marine siliciclastic deposits, and
the Balchik part, in which (sub-)littoral sandy to clayey sedi-
mentation prevailed (Figs. 1c, 2).

The following summary of basin evolution is based

mainly on Popov et al. (1986) and Popov & Kojumdgieva
(1987), who subdivided the Miocene succession into seven
formal and two informal lithostratigraphic units (Fig. 2).

The oldest Miocene rocks are biogenic limestones, 2 to

3 m thick, of Middle to Late Tarkhanian age (Karapelit Fm;
not shown in Fig. 2). They were formed in a relatively nar-
row strait along the South Dobrudzha lowlands (Fig. 1c),
which at that time connected the Eastern Paratethys with the
Fore Carpathian Basin. Within the Varna part of the depres-
sion, Tarkhanian sediments are represented by a few meters
of laminated clays near the base of Galata Formation. A ma-

Fig. 1. a – Tectonic map of Bulgaria, showing the position of the study area within the south-eastern part of the Moesian Platform (simplified
after Dabovski et al. 2002); b – Simplified geological map of the studied area (modified after Cheshitev et al. 1994a,b, 1995); c – Schematic
scheme of the paleogeography of the studied area during the Miocene (modified after Popov & Kojumdgieva 1987).

411

ORGANIC MATTER AND HYDROCARBON IN MIDDLE MIOCENE SEDIMENTS (NE BULGARIA)

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

rine regression triggered by enhanced tectonic activity inter-
rupted the connection with the Fore Carpathian Basin at the
end of the Tarkhanian.

The overlying Galata Formation is limited to the Varna part

of  the  depression  (Fig. 2).  It  is  composed  predominantly  of
mid-  to  coarse-grained  sands,  intercalated  by  frequent  clay
and  rare  limestone  and  clayey  sandstone  beds.  Ori  (2004)
investigated outcrops south of Varna and established a shal-
low  marine  environment  with  predominant  beach  sands,
tidal  deposits  and  deltaic  channels.  According  to  Popov  &
Kojumdgieva (1987), the southern part of the depression be-
came dry land at the end of the Chokrakian shifting the area of
maximum sedimentation northwards. North of Varna, deposi-
tion of the sediments of the Galata Formation continued until
Early Volhynian time (Fig. 2), but the quantity of siliciclastic
material  was  significantly  reduced  (Popov  &  Kojumdgieva
1987). At the same time up to 85 m of white-greenish calcare-
ous clays, irregularly alternating with marls, and oolitic, bio-

genic, or bioclastic limestones were deposited in the Balchik
part of the depression (Clay-limestone formation; Fig. 2). The
rocks are known from few wells and poorly studied.

The overlying Euxinograd Formation (Fig. 2) is made up

of finely laminated, diatomaceous pelites with frequent thin
(2—20 mm) coquina intercalations composed of well-stratified
shell  fragments.  Koleva-Rekalova  (1997)  also  reported  the
presence of lenses, rich in plant remains. Thin interlayers of
sandstone  or  limestone  are  very  rare.  Some  rocks  were  de-
scribed as diatomite (Popov & Kojumdgieva 1987). However,
Koleva-Rekalova  (1997)  showed  that  opal  sponge  spicules
prevail  over  diatom  frustules.  Studies  performed  by  Koleva-
Rekalova (1997, 1998), indicate deposition in a sublittoral low
energy environment, influenced by surface water currents, in-
troducing  significant  amounts  of  bioclastic  material  to  the
deeper zones of the depression. Based on the abundant fauna
Kojumdgieva & Dikova (1978), estimate a water depth of less
than  100 m,  which  is  consistent  with  the  predominance  of

Fig. 2. Lithological columns of the Miocene suc-
cession in the Varna-Balchik Depression (modified
after Popov & Kojumdgieva 1987). Correlation of
regional stages after Steininger & Wessely (2000).

412

ZDRAVKOV, BECHTEL, ĆORIĆ and SACHSENHOFER

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

benthic  foraminifera  (Darakchieva  1989)  and  with  high
amounts  of  sponge  spicules  (Koleva-Rekalova  1998).  The
thickness  of  the  Euxinograd  Formation  varies  considerably
and exceeds 100 m in the central part of the depression north
of Varna (Popov & Kojumdgieva 1987). North and west of
Balchik, the sand-limestone formation splits the Euxinograd
Formation into a lower and an upper unit (Fig. 2 – Popov &
Kojumdgieva  1987)  and  reduces  its  thickness  significantly.
Increased  terrigenous  input  from  the  north  (sand-limestone
fm)  and  south  (Franga  Fm),  was  probably  a  result  of  late
Volhynian  uplift  of  surrounding  areas  (Fig. 2  –  Popov  &
Kojumdgieva  1987),  limited  the  area  of  deposition  of  the
Euxinograd Formation to the central basin between Balchik
and  Albena  resort.  This  situation  lasted  until  the  Early
Bessarabian,  when  decreased  terrigenous  input  allowed  ex-
pansion of the area of clay deposition to the northern part of
the depression.

After the Middle Bessarabian water salinity increased rap-

idly resulting in carbonate sedimentation. The area of the ba-
sin gradually diminished and was generally restricted to the
Balchik  part  of  the  depression.  Koleva-Rekalova  (1994,
1997,  1998);  Koleva-Rekalova  &  Darakchieva  (2002)  pro-
vide a comprehensive overview of the Sarmatian rocks. Po-
rous  bioclastic  to  oolitic  limestones  with  micritic  cement
(Odarci Fm; Fig. 2) were deposited in shallow environments,
followed by massive aragonitic sediments that in upper part
show seasonal lamination (Topola Fm; Fig. 2). According to
Koleva-Rekalova (1998) these rock formed due to chemical
precipitation of needle-like aragonite crystals in shallow wa-
ter  under  specific  conditions.  The  youngest  Miocene  rocks
are  micritic  limestones  with  varying  abundance  of  bivalve
shells (Karvuna Fm; Fig. 2).

Methods

The present study is based on 47 core samples from well

C-180a drilled during the 1980s. The aim of the well was to
study  the  Lower  Cretaceous  (Valanginian)  carbonate  succes-
sion, the main aquifer in north-eastern Bulgaria. Unfortunately
there is no information on Miocene rocks. Therefore, we use
geophysical  data  from  well  C-180,  located  about  20 m  north
of C-180a, to reconstruct the lithology of the missing parts of
the studied core. Twelve additional samples were taken from
an  outcrop  near  Albena  resort  (Fig. 1b),  representing  about
5 m from the upper part of the Euxinograd Formation.

For microscopic investigations thin slides were prepared

and studied under transmitted light using polarized micro-
scope Leica DM 2500P.

Total carbon (TC) and sulphur (S) contents were deter-

mined with an Eltra Helios C/S analyser. Total organic car-
bon (TOC) content was determined on samples pretreated
with phosphoric acid. Total inorganic carbon (TIC =TC—TOC)
contents were used to calculate calcite equivalent percentages
(calcite

= TIC

×8.34). Rock-Eval pyrolysis was performed

using a Rock-Eval 6 instrument. S

and S

(mg HC/g rock)

values were used to calculate Hydrogen index (HI = 100

×S

TOC[mg HC/g TOC]) and Production index (PI = S

/(S

+ S

);

Espitalié et al. 1977). The temperature of maximum hydro-

carbon generation (T

max

) was recorded as a maturity para-

meter.

Biogenic silica contents were measured using the method

proposed by Zolitschka (1998) and a Perkin-Elmer 3000
AAS spectrometer. The silica contents were used to calculate
the opal percentages ( = biogenic silica

×2.4).

21 samples, mostly containing TOC greater than 1 wt. %,

were  chosen  for  biomarker  analysis.  Approximately  10 g  of
each  sample  were  extracted  for  approximately  1 h  using
dichloromethane  (DCM)  in  a  Dionex ASE 200  accelerated
solvent  extractor.  Asphaltenes  were  precipitated  from  a  hex-
ane-DCM  solution  (80 : 1)  and  separated  by  centrifugation.
The  hexane-soluble  fractions  were  subdivided  into  saturated
and  aromatic  hydrocarbons  and  NSO  components  using  a
Köhnen-Willsch  MPLC  instrument  (Radke  et  al.  1980).  No
separation of the aliphatic and aromatic fractions was conducted
only in the case of sample C-20, due to the low amount of ex-
tract. This sample was analysed by GC-MS as total fraction.

The saturated and aromatic hydrocarbons fractions were

analysed by a gas chromatograph-mass spectrometer (GC-MS)
Thermo Fisher Trace Ultra, equipped with a silica capillary
column. Oven temperature was programmed from 70-300 °C
with steps of 4 °C/min, followed by an isothermal period of
15 min. Helium was used as the carrier gas. The device was
set in electron impact mode with a scan rate of 50—650 Daltons
(0.7 sec/scan).  Data  were  processed  with  Thermo-Fisher
Xcalibur  ®  v. 2.0.  Identification  of  biomarkers  is  based  on
retention  time  and  comparison  of  mass  spectra  with  pub-
lished data. The determination of absolute concentrations of
biomarkers  was  done  using  internal  standards  (deuterated
n-tetracosane  for  the  aliphatic  fraction  and  1,1’-binaphthyl
for the aromatic fraction). The concentrations were normal-
ized against the total organic carbon contents.

Carbon isotope determination of n-alkanes and isoprenoid

hydrocarbons was performed on 11 samples using a Trace GC
ultra attached to a ThermoFisher DELTA-V ir-MS via a com-
bustion interface (GC Isolink, ThermoFisher). For calibration,
a CO

standard was injected at the beginning and end of each

analysis. The GC coupled to the ir-MS was equipped with the
column described above and the temperature program was the
same as for GC-MS analysis. Isotopic composition is reported
in the

δ notation relative to the PDB standard.

Results and discussion

The lithological profile of well C-180a is plotted together

with the resistivity log in Fig. 3. Based on lithology, the Mio-
cene succession can be divided into five units, which in gen-
eral correspond to the lithostratigraphic units introduced by
Popov & Kojumdgieva (1987).

Lithology

Unit I

Unit I discordantly overlies finely laminated Oligocene

brownish clays and correlates with the Clay-limestone for-
mation from the northern part of the Varna-Balchik Depres-

413

ORGANIC MATTER AND HYDROCARBON IN MIDDLE MIOCENE SEDIMENTS (NE BULGARIA)

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

sion and with the Galata Formation from its southern part
(Popov & Kojumdgieva 1987). The unit is 87 m thick (Fig. 3)
and can be subdivided into two subunits.

Subunit Ia – is about 35 m thick and is represented by

white-greenish, very soft fine-grained glauconitic sands. The
rocks appear homogenous, but sedimentary structures within
the  non-lithified  sands  may  have  been  obliterated  during
coring.  Carbonate  contents  are  typically  low  (<3 wt. %;
Table 1),  but  reach  up  to  25 wt. %  in  the  upper  part  of  the
section. Carbonate is exclusively represented by micrite crys-
tals.  No  biogenic  carbonate  remains  were  detected.  Micro-
scopic  investigations  showed  that  the  detrital  material  is
predominantly of fine- and very fine sand-size fractions. The
mineral composition is relative uniform and is dominated by
quartz and glauconite. Rare feldspar (mostly orthoclase; rare
microcline), as well as mica flakes, can also be found. Some
samples  contain  lithic  fragments,  composed  of  polycrystal-
line  quartz,  and  fine  charcoal  fragments.  Irregular-shaped,

subangular to subrounded quartz grains are the predominant
detrital  component.  Significant  amounts  (15—30 wt. %)  of
light green glauconitic pellets with very fine sand sizes give
the rocks a pale greenish colour. Under the microscope, most
glauconite  grains  show  red-brownish  colours,  indicating
initial  oxidation.  Although  an  authigenic  origin  of  glau-
conite is possible, we assume a detrital origin, because glau-
conite-bearing  Cretaceous  and  Paleogene  rocks  exist  in  the
hinterland.

Subunit Ib – about 52 m thick (Fig. 3), is composed of

hard,  but  highly  porous  biomicritic  limestone  layers  with
randomly  orientated  intact  and  fragmented  bivalve  shells.
These rocks alternate irregularly with marlstones, up to 1 m
thick.  A  thin  layer  of  oolitic  limestone  was  observed  in  the
upper part of the subunit. Although the core representing the
lower  part  of  subunit Ib  is  missing,  a  gradual  upward  in-
crease  in  resistivity  suggests  a  continuous  transition  from
subunits Ia to Ib.

Popov et al. (1986) give a similar description for the rocks

from  the  Clay-limestone  formation  in  two  other  drill  holes
from  the  Balchik  area.  Considering  the  above  data,  we  as-
sume that deposition of subunit Ia commenced in a sand and
mud  dominated  shallow  marine  environment,  characterized
by high terrigenous input. The small grain size indicates for-
mation  in  a  relatively  low  energy  environment,  probably
within the inner shelf below the wave-dominated zone. Ac-
cording to Kojumdgieva (1965), the Chokrakian was a time
of reduced salinity. Very low sulphur contents in sediments
of unit Ia may support this interpretation.

During the later stages of subunit Ia sediment deposition, a

gradual  decrease  of  the  terrigenous  input  enabled  increased
carbonate deposition. Apart from this, the sedimentary envi-
ronment  most  probably  did  not  change  significantly.  The
dominance of biomicritic limestone with random orientation
of  the  fossils  in  subunit Ib  argues  for  deposition  in  shallow
marine  environments  with  bottom  currents  and/or  waves.
Oolitic  limestone  interbeds  in  the  upper  part  of  the  subunit
indicate  a  change  to  a  very  shallow,  high  energy  environ-
ment (Tucker & Wright 1990).

Unit II

Unit II, about 47 m thick, follows above unit I with a sharp

boundary. Similar to unit I, two subunits can be distinguished
(Fig. 3).

Subunit IIa – is composed of 13-m-thick sandy limestone

in  transition  to  fine  sand  with  micrite  matrix.  The  rocks  are
white to white-grey, very soft and friable. The carbonate con-
tent varies significantly between 21 and 80 wt. % and is rela-
tive  low  in  the  middle  part  of  the  subunit  (Table 1,  Fig. 4).
Carbonates  are  almost  exclusively  represented  by  micrite
crystals. With the exception of few foraminifera, detected by
light microscopy, no other organic remains were detected.

The siliciclastic component is composed of fine and very

fine sand-sized quartz grains and rare feldspar and glauco-
nite grains. The clastic fraction is further characterized by
the presence of detrital mudstone fragments, often including
small charcoal fragments. This shows that at least part of the
organic matter is redeposited.

Fig.  3.  Lithologi-
cal profile of well
C-180a,  together
with  geophysical
resistivity log.

414

ZDRAVKOV, BECHTEL, ĆORIĆ and SACHSENHOFER

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

Table 1: Depth from surface, sulphur (S), total organic carbon (TOC), total inorganic carbon (TIC), calcite equivalent (eq.) percentages, biogenic
silica (BSi), TOC/S ratio, Rock Eval parameters (S

, S

, T

max

), hydrogen index (HI), production index (PI), soluble organic matter yield (SOM),

and relative proportion of saturated (Sat. HC) and aromatic (Aro. HC) hydrocarbons, polar compounds (NSO) and asphaltenes (Asph.).

Sample Depth Unit

S TOC TIC Calcite eq. BSi

TOC/S

Tmax

HI PI

SOM

(mg/g TOC)

Sat. HC Aro. HC NSO Asph.

(wt. %)

mg HC/g

rock

(°C)

(wt. %, SOM)

Drillhole C-180a
C-1

107.0

0.2 1.6 6.6

54.9 2.4

6.7 0.2 4.2 423 254 0.04

19 16

C-2

107.4 0.7

1.6

6.4

53.0 2.6

2.4 0.2 3.7 424 229 0.04

C-3

107.9 0.0

0.7

9.9

82.3 1.2

93.6 0.1 1.6 429 210 0.05

29 18

C-4

108.2 1.0

2.6

4.2

35.3 6.5

2.6 0.4 7.9 422 301 0.05

C-5

108.5 1.4

4.3

4.0

33.1 1.9

3.0 0.5 8.6 423 201 0.05

25 22

C-6

108.8 1.5

3.7

3.4

28.1 3.9

2.5 0.4 7.5 422 205 0.05

C-7

109.0 1.0

3.2

5.2

43.4 2.4

3.1 0.3 6.5 422 203 0.04

21 12

C-8

109.3 1.2

2.9

4.7

39.4 1.7

2.4 0.3 5.2 424 176 0.05

C-9

109.5 0.4

1.8

7.2

60.0 0.9

4.8 0.1 3.2 423 179 0.04

18 16

C-10

109.8 0.4

1.2

7.3

60.6 1.0

3.4 0.1 2.8 425 227 0.04

C-11

110.2 0.5

1.3

6.6

55.2 2.4

2.8 0.1 2.2 425 174 0.04

35 34

C-12

110.6 0.3

1.2

6.2

51.6 2.4

4.8 0.1 3.3 422 272 0.03

26 18

C-13

110.7 0.2

0.6

8.0

66.8 1.0

3.3 0.1 1.0 428 172 0.06

C-14

111.5 0.7

1.1

6.0

49.7 1.9

1.7 0.2 3.8 425 333 0.04

35 14

C-15

112.0 0.5

1.3

5.1

42.5 1.9

2.5 0.1 1.7 426 129 0.05

30 16

C-16

113.0 0.0

0.6

8.6

71.6 1.0

49.8 0.1 1.3 429 213 0.05

51 42

C-17

113.7 0.0

0.6

8.8

73.4 2.0

80.9 0.1 1.2 429 194 0.05

37 12

C-18

114.0 0.0

0.6

8.3

69.0 1.7

94.6 0.1 0.9 428 136 0.06

C-19

114.2 0.0

0.6

7.6

63.1 1.9 184.7 0.0 0.6 433 96

0.06

C-20

114.5 0.0

0.2

10.4

86.4 0.9 146.7 –

–

41 –

–

C-21

148.7

Uni

t IIa

0.0 1.7 8.1

67.7 0.5

63.1 0.2 3.8 420 223 0.06

16 31

C-23

149.4 0.0

0.7

5.6

47.0 9.0

37.3 0.2 2.2 419 318 0.06

C-24

149.8 0.0

0.3

7.7

63.8 5.1

45.9 0.1 1.0 419 309 0.07

C-25

150.3 0.0

0.8

8.0

66.3 2.1

48.4 0.2 2.5 419 318 0.07

C-26

152.0 0.5

1.1

4.5

37.8 3.3

2.2 0.3 4.3 416 397 0.07

21 9

C-27

153.0 0.0

0.9

4.6

38.3 11.8

38.0 0.3 3.4 415 359 0.08

C-28

153.9 0.0

1.0

7.8

64.9 14.3

47.2 0.3 3.8 421 360 0.07

20 24

C-29

154.7 0.1

0.8

2.9

24.4 4.5

6.2 0.2 3.4 418 427 0.06

C-30

155.2 0.0

0.5

6.6

55.0 8.7

27.3 0.1 1.4 418 287 0.07

C-22

155.7 0.9

2.8

2.5

20.9 2.4

3.1 0.5 7.3 419 257 0.06

14 31

C-31

157.9 0.0

1.0

3.7

31.0 22.1

40.4 0.4 4.8 419 464 0.08

40 18

C-32

158.6 0.0

0.9

4.2

35.1 11.1

46.3 0.3 4.4 422 501 0.06

C-34

159.7 0.0

0.2

6.7

55.5 1.5 260.2 0.0 0.3 423 186 0.08

C-35

160.5 0.0

0.7

8.4

69.7 1.5

42.9 0.1 1.4 420 209 0.06

18 22

C-36

161.5 0.0

0.6

9.6

80.2 1.3

43.2 0.1 1.6 423 248 0.07

C-37

214.3

Uni

t Ia

0.3 0.1 0.3

2.3 –

0.2 –

–

C-38

215.3 0.1

0.1

2.3

19.5 –

0.6 –

–

C-39

216.4 0.2

0.1

2.4

19.6 –

0.4 –

–

C-40

219.2 0.0

0.2

3.1

25.5 –

33.9 –

–

C-41

221.0 0.0

0.0

0.9

7.1 –

11.0 –

–

C-42

230.0 0.0

0.1

0.5 –

18.7 –

–

C-43

231.0 0.0

0.0

0.2 –

17.7 –

–

C-44

232.0 0.0

0.0

0.4 –

11.2 –

–

C-45

236.0 0.1

0.1

0.8 –

0.5 –

–

C-46

238.0 0.0

0.0

0.2

1.5 –

13.4 –

–

C-47

240.0 0.0

0.1

0.9 –

16.4 –

–

C-48

241.0 0.0

0.1

0.3

2.6 –

44.2 –

–

Outcrop near Albena resort*

A-1

3.7

Uni

t III

0.0 0.1 0.2

1.4 8.2

1.9 – –

–

A-2

4.1 1.1

0.9

1.5

12.2 8.5

0.8 0.1 1.5 411 166 0.06

A-3

4.4 0.8

1.1

1.8

14.7 12.8

1.4 0.2 2.4 414 216 0.06

21 11

A-4

4.8 1.0

1.2

1.9

16.2 1.4

1.2 0.2 2.5 417 213 0.06

A-5

5.2 0.4

1.0

2.2

18.4 11.6

2.7 0.1 2.0 417 200 0.05

A-6

5.6 0.3

0.7

3.0

25.0 7.3

2.2 0.1 1.3 413 171 0.06

A-7

6.0 0.6

1.1

2.5

21.1 10.6

1.8 0.1 2.4 412 207 0.02

15 12

A-8

6.4 0.4

0.7

3.4

28.2 3.7

1.7 0.2 0.6 419 77

0.22

A-9

6.7 0.6

1.1

2.8

23.3 8.6

1.8 0.1 2.4 419 213 0.04

A-10

7.1 0.5

1.0

3.5

29.1 9.9

2.0 0.1 2.5 415 256 0.05

A-11

7.6 0.5

1.3

4.1

33.8 5.5

2.5 0.1 2.2 415 175 0.05

19 12

A-12

8.1 0.5

1.1

2.7

22.3 1.6

2.1 0.1 2.4 423 220 0.04

* — Sampling position in meters above sea level.

415

ORGANIC MATTER AND HYDROCARBON IN MIDDLE MIOCENE SEDIMENTS (NE BULGARIA)

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

Fig. 4. Vertical distribution of total organic carbon (TOC), sulphur, calcite equivalent percentages, biogenic silica (BSi), hydrogen index (HI),
relative proportions of n-alkanes, CPI, pristane/phytane ratio, sterane and hopane concentrations in core samples from subunit IIa and
unit III, and outcrop samples.

416

ZDRAVKOV, BECHTEL, ĆORIĆ and SACHSENHOFER

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

Rocks from subunit IIa contain varying amounts of bio-

genic silica derived from diatom frustules (Table 1). Diatom
remains  were  previously  detected  mostly  within  clays  of
the  Euxinograd  Formation  (Popov  &  Kojumdgieva  1987;
Temniskova-Topalova  1990).  This  suggests  that  subunit IIa
is  coeval  with  the  lower  part  of  this  formation.  Biogenic
silica  contents  reach  a  maximum  of  22 wt. %  about  4 m
above  the  base  of  the  subunit  (C-31)  and  decrease  upwards
(Fig. 4,  Table 1).  However,  the  biogenic  silica  content  may
significantly under estimated diatom contents. For example,
microscopic  inspection  revealed  that  sample  C-23  (9 wt. %
silica) contains significantly more than 50 % of diatoms. The
abundance of diatoms reflects a high nutrient supply and dis-
solved  silica  (DeMaster  2003).  High  nutrient  supply  is  also
supported  by  the  frequent  occurrence  of  calcareous  nanno-
plankton  assemblage  rich  in  Braarudosphaera  bigelowii
(Gran  &  Braarud  1935)  Deflandre,  1947  in  sample  C-31,
which also argues for reduced salinity and cold water envi-
ronment.  High  nutrient  and  silica  supply  may  be  related  to
rhyolitic volcanic activity, indicated by tephra layers within
the Euxinograd Formation (Milakovska et al. 2001). The oc-
currence of Cyclicargolithus floridanus (Roth & Hay in Hay
et  al.  1967)  Bukry,  1971  shows  that  sample  C-23  is  not
younger than 13.33 Ma (Lourens et al. 2004).

Subunit IIb – is represented by white to grey-white

biomicritic limestone, similar to that in subunit Ib. Although
significant parts of the core are missing, the resistivity log
suggests the presence of limestone within the missing inter-
vals. The similarity of the lithologies of units I and II suggest
that the rocks formed in similar sedimentary environments.

Unit III

Unit III in well C-180a is about 14 m thick and represents

the  upper  part  of  the  Euxinograd  Formation  (Figs. 2, 3).  Its
characteristics  are  similar  to  the  5-m-interval  studied  in  the
outcrop section. At both sites unit III is composed of a mo-
notonous  succession  of  fine-laminated  clays  and  detritus  of
bivalve  shells.  With  the  exception  of  a  few  samples,  the

Fig. 5. Graphs of TOC versus calcite equivalent percentages.

rocks contain no detrital sand-sized grains. Needle-shaped
siliceous sponge spicules are extremely abundant in the up-
per part of the unit (samples C-10 to C-1) where biogenic
silica contents reach up to 6.5 wt. % (Table 1).

The presence of sponge spicules proves deposition in a low

energy  sedimentary  environment  (Krautter  1998).  A  low  en-
ergy  environment  and  oxygen-depleted  conditions  are  also
supported  by  the  lamination  of  the  rocks  (see  also
Kojumdgieva  &  Dikova  1978;  Koleva-Rekalova  1998).  The
presence  of  ascidian  spicules  indicates  shallow  water  condi-
tions  and  enrichment  in  the  nutrient  supply  (Toledo  et  al.
2007). Whereas sponge spicules are abundant in well C-180a,
diatoms  are  more  frequent  in  the  outcrop  samples,  maybe
due to different distance from the shoreline.

Redeposited coaly organic material is present, as in sub-

unit IIa.

Units IV and V

Units IV and V correlate with the Upper Miocene Topola

and  Karvuna  Formations  (Popov  &  Kojumdgieva  1987),
respectively.  They  are  represented  mainly  by  white  mas-
sive  and  fine  laminated  aragonitic  sediments  (Topola  Fm;
Fig. 2 – Koleva-Rekalova 1994) and biomicritic limestones
(Karvuna Fm; Fig. 2). Since the rocks do not contain organic
matter, they are not considered in this study.

TOC and Rock-Eval parameters

Total organic carbon (TOC) contents are listed in Table 1

and  are  plotted  versus  depth  in  Fig. 4.  TOC  contents  in
subunit Ia  are  very  low  ( < 0.2 wt. %;  Table 1)  and  vary  be-
tween 0.2 and 4.3 wt. % in the rest of the samples. The high-
est TOC contents (1.8—4.3 wt. %) were detected in the upper
part  of  unit III  from  well  C-180a  (Table 1,  Fig. 4).  Single
samples with TOC contents up to 2.8 wt. % also exist in the
middle and upper part of subunit IIa (Table 1, Fig. 4).

Fig. 5 shows the relationship between TOC and calcite eq.

percentages. Whereas there is no linear relation for units Ia

417

ORGANIC MATTER AND HYDROCARBON IN MIDDLE MIOCENE SEDIMENTS (NE BULGARIA)

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

and IIa, and the studied outcrop (unit III), a strong negative
relation  (correlation  coefficient:  —0.86)  exists  for  samples
from  unit III  from  well  C-180a.  Ricken  (1991)  points  out
that a negative correlation is indicative for roughly constant
organic  matter  production  and  dilution  by  varying  amounts
of calcite. Calcite in this unit is mostly derived from calcare-
ous bivalve shell fragments.

Because of very low TOC contents, TOC/S ratios have not

been  calculated  for  samples  from  subunit Ia.  For  samples
from  units IIa  and  III  TOC/S  ratios  range  between  1.7  and
260.2.  In  general,  the  ratios  are  higher  in  subunit IIa  (me-
dian:  42.9)  than  in  unit III  (3.4).  High  values  in  subunit IIa
are  consistent  with  a  salinity-reduced  environment.  The
TOC/S trend in unit III suggests an upward increase of salin-
ity  to  normal  marine  conditions  during  deposition  of  the  up-
permost part of the Euxinograd Formation. Anoxic conditions
indicated by TOC/S ratios below 2.8 (Berner 1984; Raiswell
& Berner 1986) are not reflected by the data.

Hydrogen index (HI) values are summarized in Table 1.

Their vertical distribution is shown in Fig. 4. Peters (1986)
points out that pyrolysis experiments usually give erroneous
results in samples with low TOC contents. For that reason HI
was not calculated for samples, containing less than 0.2 wt. %
TOC. These are mostly from subunit Ia, but also samples
C-20 and A-1 from unit III. HI values vary significantly be-
tween 77 and 501 mg HC/g TOC.

Surprisingly, the sandy subunit IIa exhibits higher HI val-

ues  (Fig. 6a;  average:  324 mg HC/g TOC),  than  the  clayey
unit III  (205  and  192  for  core  and  outcrop  samples,  respec-
tively).  This  is  also  reflected  in  plots  of  S

versus TOC

(Langford & Blanc-Valleron 1990) and HI versus T

max

(Fig. 6b – Espitalié et al. 1984). Whereas kerogen types II-III
prevails  in  unit III,  kerogen  type II  predominates  in  sub-
unit IIa. Because biomarker data (see below) argue against a
higher  contribution  of  landplants  to  the  organic  matter  in
unit III,  the  difference  might  reflect  enhanced  microbial  al-
teration of the organic matter in unit III.

Hydrocarbon potential

Peters (1986) proposed a classification scheme for source

rocks based on TOC, S

and HI values. The respective average

values for subunit IIa (TOC : 0.9 wt. %; S

: 3.0 mg HC/g rock;

HI : 324 mg HC/g TOC) show that it holds a fair to good hy-
drocarbon generative potential for oil. The upper 5 m of unit III
(TOC:1.7 wt. %; S

:3.5 mg HC/g rock; HI:205 mg HC/g TOC)

show that this thin interval holds a good to fair potential with
the possibility to generate gas and minor oil.

The low maturity and the relatively low thickness of both

diatom-bearing  intervals,  which  are  considered  equivalents
of  the  Euxinograd  Formation,  prevent  significant  hydrocar-
bon  generation  in  onshore  areas.  However,  both,  thickness
and maturity may increase towards the east in offshore areas.

Molecular composition of the hydrocarbons

The normalized yields of the soluble organic matter (SOM)

are presented in Table 1 together with the contents of satur-
ated  and  aromatic  hydrocarbons,  hetero-compounds  (NSO)
and  asphaltenes.  The  contents  of  SOM  are  generally  low
(14—51 mg/g TOC), which is in accordance with a low matu-
rity of the organic matter. In all samples, the SOM is mainly
represented by polar compounds (37—84 wt. %; Table 1). It is
noticeable,  that  most  borehole  samples  show  increased  con-
tents of saturated hydrocarbons and are depleted in aromatics
and  asphaltenes,  in  relation  to  outcrop  samples.  This  finding
might be an indication of the presence of free hydrocarbons in
these samples. However, very low PI values (Table 1) are not
consistent with the presence of mobile products.

n-Alkanes

The total ion current (TIC) chromatograms of the saturated

hydrocarbon fractions of eight samples are shown in Fig. 7.
The vertical trends of the n-alkanes and CPI (Carbon Prefer-
ence Index after Bray & Evans 1961) in units IIa and III in
well C-180a, are presented in Fig. 4.

Fig. 6. Plots of S2 versus TOC (a) and HI versus T

max

(b). Dashed

lines in Fig. 6a represent HI values of 200 mgHC/gTOC (kerogen
type III) and 700 mgHC/gTOC (kerogen type I).

418

ZDRAVKOV, BECHTEL, ĆORIĆ and SACHSENHOFER

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

All samples are characterized by the presence of n-alkanes

in the range of n-C

to n-C

(Fig. 7). However, their distri-

butions vary justifying the definition of three groups.

(i) Most samples from subunit IIa and six samples from

unit III are characterized by a predominance of short-chain
n-alkanes (n-C

15—20

; Table 2, Fig. 7) and contain minor amounts

of mid- and long-chain hydrocarbons. Typically, the domi-
nant n-alkane is n-C

, accompanied by appreciable amounts

of n-C

and n-C

, with no odd-over-even predominance.

(ii) Outcrop samples from unit III show distributions domi-

nated by mid- (n-C

21—25

) and long-chain n-alkanes (n-C

27—33

;

e.g. sample A-7 in Fig. 7), with maximum concentrations of
n-C

or n-C

. The CPI is moderately high (1.7—2.2; Table 2),

and short-chain hydrocarbons are usually absent.

(iii) Another six borehole samples from unit III are domi-

nated either by n-C

20—22

, with a maximum at n-C

(e.g. C-5,

C-9) and a CPI of 1.2 to 1.7, or by n-alkanes in the range of
n-C

23—33

with a low CPI (1.1—1.3; e.g. C-11; Fig. 7).

Short-chain n-alkanes originate mostly from algae and mi-

croorganisms (Blumer et al. 1971; Cranwell 1977). Han et al.
(1968) and Han & Calvin (1969) established a marked pre-
dominance of n-C

in blue-green algae, and appreciable

amounts of n-C

and n-C

in photosynthetic and non-photo-

synthetic bacteria. Considering the predominance of n-C

, a

significant contribution from phytoplankton and microorgan-
isms can be suggested for samples with a type (i) distribution.

Mid-chain n-alkanes (n-C

21—25

) are typical mostly for sub-

merged aquatic plants and some moss species (Cranwell

Fig. 7. Gas chromatograms of the saturated hydrocarbon fraction (Std. = standard). The numbers correspond to n-alkanes with the respective
carbon atoms. Asterisk numbers correspond to iso-alkanes with the respective carbon atoms.

419

ORGANIC MATTER AND HYDROCARBON IN MIDDLE MIOCENE SEDIMENTS (NE BULGARIA)

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

Table 2: Molecular composition of the organic matter.

-C

15—20

/n-C

15—33

= ; n-C

21—25

/n-C

15—33

= ; n-C

27—31

/n-C

15—33

= .

CPI — Carbon Preference Index (according to Bray & Evans 1961),

CPI = 0.5*( + ).

Pri – Pristane, Phy – Phytane;

C-27/steranes = ; C-28/steranes = ; C-29/steranes = .

Ts – 18

α-22,29,30-trisnorneohopane, Tm – 17α-22,29,30-trisnorhopane, MN – methyl naphtalenes, DMN – dimethyl naphtalenes,

TMN – trimethyl naphtalenes, P – phenanthrene, MPh – methyl phenanthrenes, PAH – polyaromatic hydrocarbons (fluoranthrene
+ pyrene + benzo[a]pyrene + chrysene).

Aryl Isoprenoids =

di-MTTC/tri-MTTC = .

i=15

(n-Ci)

i=15

(n-Ci)

i=21

(n-Ci)

i=15

(n-Ci)

i=27

(n-Ci)

i=15

(n-Ci)

(n-C

+n-C

)

(n-C

+n-C

)

(n-C

+n-C

)

(n-C

+n-C

)

1977;  Ficken  2000).  However,  Han  &  Calvin  (1969)  estab-
lished  that  mid-chain  n-alkanes  are  typical  components  of
bacterial lipids, especially in sulphate-reducing bacteria, thus
indicating  that  microorganisms  could  also  be  a  source  of
these hydrocarbons.

Long-chain n-alkanes with a high CPI are characteristic of

waxes  from  vascular  plants  (Eglinton  &  Hamilton  1967).
This suggests a major contribution of land plants for outcrop
samples with a type (ii) distribution and a minor contribution
for  type  (i)  samples.  However,  other  possible  sources  for
long-chain  n-alkanes  have  also  been  recognized.  These  in-
clude different bacterial communities, which produce mostly
hydrocarbons with lower odd-over-even predominance (Albro
&  Dittmer  1967;  Han  &  Calvin  1969;  Johnson  &  Calder

1973;  Volkman  et  al.  1980,  1983;  Lichtfouse  et  al.  1994).
Furthermore bacterial degradation of organic matter is known
to reduce the CPI (Allen et al. 1971; Johnson & Calder 1973;
Gagosian  et  al.  1983;  Meyers  1994;  Volkman  &  Tanoue
2002).  Considering  the  partly  high  amount  of  long-chain
n-alkanes with a low CPI in samples with a type (iii) distri-
bution,  a  significant  contribution  of  bacterial  sources  to  the
long-chain  n-alkanes  can  be  assumed.  As  seen  in  Table 3,
the  isotopic  fractionation  of  the  long-  and  short-chain  n-al-
kanes do not differ significantly, which might be due to their
similar origin. Nevertheless, the higher plant contribution to
the long-chain alkanes is also noticeable by the slight deple-
tion  in

C ( ~ 0.5 ‰) of these compounds. On the other

hand, redeposition of terrestrial organic matter might also

Σ αααC

(20S+20R)

i=27

αααC

(20S+20R)

ΣαααC

(20S+20R)

i=27

αααC

(20S+20R)

ΣαααC

(20S+20R)

i=27

αααC

(20S+20R)

i=13

trimetyl (2,3,6)-alkylbenzol.

dimethylated 2-methyl-2-(4´,8´,12´-trimethyltridecyl)chroman

trimethylated 2-methyl-2-(4´,8´,12´-trimethyltridecyl)chroman

n-Alkanes

Isoprenoids

Steranes Hopanespanes

Naphthalenes

(

)

-C

15–20

-C

15–33

-C

21–25

-C

15–33

-C

26–31

-C

15–33

(

)

Pr/Ph

(

)

/Σ

era

/Σ

era

/Σ

era

(

)

S+22

-31-

p-1

)-e

(µ

g T

erp

oid

(µ

g T

MN (µ

g T

DMN (µ

g T

(

)

(

)

(

)

ryl

sop

ren

oid

(µ

g T

Chr

(µ

g T

TTC

ri-

TTC

C-1

102.8 0.5 0.2 0.2 2.4 23.8 1.3 0.6 0.4 0.1 0.5

7.0 0.23 0.43 0.5 0.0 0.1 1.5 1.1 0.6 0.6 0.8 0.07 0.3

C-3

208.3 0.6 0.2 0.1 2.5 54.1 1.0 0.1 0.3 0.2 0.5

3.0 0.36

– 0.0 0.0

0.0 0.7 1.3 1.6 0.6 1.5 0.02 0.3

C-5

747.5 0.1 0.7 0.1 1.2 33.6 0.9 1.0 0.5 0.2 0.4

2.9

–

0.63 0.6 0.0 0.1 5.2 4.2 1.3 1.9 3.1 0.26 0.6

C-7

226.9 0.2 0.4 0.2 1.3 21.5 0.9 0.4 0.4 0.3 0.3

1.5

–

0.38 0.6 0.3 0.1 0.8 0.8 1.0 0.5 1.5 0.19 0.7

C-9

188.8 0.4 0.4 0.1 1.7 36.0 1.0 1.3 0.5 0.3 0.3

3.8 0.23 0.54 0.2 0.0 0.2 1.6 2.0 2.8 1.4 1.7 0.24 0.2

C-11

795.5 0.2 0.4 0.2 1.2 79.8 0.8 1.3 0.4 0.4 0.3

5.2

–

1.9 0.9 0.4 4.1 4.5 2.2 1.1 5.3 0.01 0.7

C-12

460.8 0.2 0.6 0.1 1.3 50.6 1.3 1.3 0.5 0.2 0.3

2.9 0.25

–

0.0 0.0 0.1 0.7 0.8 0.6 0.4 1.9 0.04 0.4

C-14

314.5 0.5 0.3 0.2 1.8 60.1 0.9 5.6 0.4 0.1 0.5 11.8 0.36 0.47 0.4 0.8 0.0 0.7 1.2 0.6 0.4 1.2 0.08 0.3

C-15

205.2 0.6 0.2 0.1 1.9 41.1 0.7 0.3 0.7 0.2 0.2

6.6 0.15 0.51 0.2 1.2 0.2 1.2 0.9 0.5 0.4 0.9 0.02 0.1

C-16 3132.8 0.1 0.5 0.3 1.1 70.1 0.8 0.3 0.4 0.1 0.5

–

0.0 0.0 0.7 1.2 1.0 0.7 1.5 0.09 0.1

C-17

253.2 0.6 0.2 0.1 1.8 68.9 0.9 0.8 0.5 0.3 0.2

2.3 0.25 0.59 0.2 1.3 0.5 4.4 3.2 1.3 1.3 1.9 0.03 0.1

C-20

469.8 0.3 0.3 0.3 1.4 84.3 1.0

–

8.7

–

C-21

71.5 0.3 0.4 0.2 1.2 12.2 0.9 0.4 0.4 0.2 0.4

7.4 0.80 0.65 0.8 0.8 0.0 0.0 0.0 0.6 0.4 0.0 0.00

–

C-26

368.2 0.5 0.3 0.1 1.4 40.2 0.6 0.4 0.4 0.2 0.4

3.8 0.50 0.57 1.2 1.2 0.0 0.3 0.5 0.5 1.0 0.0 0.00

–

C-28

42.2 0.5 0.2 0.1 2.1 13.2 0.7 0.0 0.4 0.2 0.4

2.4 0.50 0.63 0.2 0.2 0.0 0.1 0.1 0.2 0.2 0.0 0.00

–

C-22

27.5 0.6 0.2 0.1 2.2

8.0 0.8 0.4 0.5 0.2 0.4

0.8 0.34 0.66 0.2 0.2 0.0 0.1 0.1 0.1 0.4 0.0 0.00

–

C-31

155.2 0.4 0.2 0.2 2.0 36.1 0.7 0.3 0.2 0.2 0.6

2.4 0.54 0.65 0.2 0.2 0.1 0.7 0.8 0.6 1.7 0.0 0.00

–

C-35

86.3 0.4 0.3 0.2 1.7 11.9 0.5 0.6 0.4 0.2 0.4

1.8 0.66 0.66 0.0 0.0 0.0 0.2 0.6 0.1 4.0 0.0 0.00

–

A-3

164.1 0.1 0.4 0.4 2.0

2.8 0.6 0.4 0.4 0.2 0.4

1.5 0.57 0.57 0.2 0.2 0.2 0.0 0.0 0.1 0.1 0.0 0.01 0.2

A-7

163.8 0.0 0.4 0.4 2.2 1.3 0.6 0.1 0.4 0.2 0.4

2.0 0.36 0.41 0.1 0.1 0.2 0.0 0.0 0.1 0.2 0.0 0.01 0.2

A-11

154.5 0.0 0.4 0.4 1.7 0.9 0.7 0.2 0.4 0.2 0.4

2.7 0.54 0.51 1.3 1.3 0.2 0.0 0.6 1.0 4.7 0.0 0.07 0.1

420

ZDRAVKOV, BECHTEL, ĆORIĆ and SACHSENHOFER

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

represent a substantial factor for the low CPI values, ob-
served in the studied samples.

Isoprenoids

In the studied samples, the isoprenoids are represented by

a series of compounds in the C

14—20

range (Table 2). Among

them, pristane and phytane are the most important, because
of their ubiquity in sediments and oils of different maturity
(Brooks et al. 1969). The relative proportions of isoprenoids
depend  on  the  Eh  setting  of  the  environment,  since  anoxic
conditions  favour  the  formation  of  phytane  and  more  oxic
conditions favour the formation of pristane (e.g. Volkman &
Maxwell  1986).  Hence,  the  pristane/phytane  (Pr/Ph)  ratio
can be used to evaluate the Eh potential of the depositional
environment (Didyk et al. 1978).

However, because pristane is formed during maturation of

the organic matter, the Pr/Ph ratio is also influenced by ther-
mal maturity (Goossens et al. 1984; Volkman & Maxwell
1986; Hughes et al. 1995; Koopmans et al. 1999). Moreover,
pristane may be formed from the isoprenoidal side-chain of to-
copherols (Goossens et al. 1984; ten Haven et al. 1987), phy-
toplankton and bacteria (Han et al. 1968; Han & Calvin 1969;
Volkman 1988), and phytane may be formed from bacterial
lipids (e.g. Volkman & Maxwell 1986; Volkman 1988).

Minor contributions from additional sources to both pris-

tane and phytane cannot be completely discarded. However,
since  the  low  maturity  of  the  rocks  suppresses  a  significant
contribution to pristane concentrations due to thermal matu-
rity,  the  Pr/Ph  ratio  is  considered  an  indicator  of  the  oxic/
dysoxic  depositional  conditions.  Furthermore,  the  isotopic
ratios of both pristane and phytane do not differ significantly
(Table 3),  which  might  be  an  indication  for  their  common
origin from chlorophyll-a. The Pr/Ph ratio is 1.3 for samples
C-12  and  C-1  from  the  middle  and  upper  parts  of  unit III
(Table 2, Fig. 4) indicating dysoxic conditions. The Pr/Ph of
all remaining samples ranges between 0.5 and 1.0 (Table 2),

Sample

A-7 C-1 C-5 C-7 C-11 C-14 C-16 C-21 C-22 C-26 C-35

n-C

–28.3 –28.2 –28.7 –29.1 –29.3

n-C

–28.6

–28.3

–28.8

–29.4

–29.1

–29.5

–29.0

–29.5

–28.6

n-C

–28.5 –28.7 –28.7 –29.5 –29.4 –29.9 –29.6 –29.4 –29.1 –29.1

Pristane

–30.1 –29.8 –29.8 –30.2 –30.1 –30.4 –29.9 –30.2 –30.4 –29.9

n-C

–29.0 –28.7 –28.7 –29.5 –29.3 –29.4 –29.1 –29.3 –28.5 –29.4

Phytane

–29.9 –29.8 –30.1 –30.7 –30.5 –30.2 –30.1 –30.2 –30.1 –30.2

n-C

–28.5

–28.6

–28.5

–28.4

–29.0

–28.7

–28.8

–28.7

–29.2

n-C

–28.1 –28.3 –28.7 –28.5 –28.0 –28.3

–29.2 –27.9 –29.0

n-C

–27.9 –28.2 –28.1 –28.8 –27.7

–28.6 –28.5

n-C

–28.3

–28.1 –27.7 –27.2 –28.1 –27.4

–27.6 –28.0

n-C

–27.5

–28.3 –28.1 –27.0 –28.3 –27.6

–28.5 –27.9

n-C

–28.0

–27.5

–26.9

–26.8

–27.8

–26.9

–27.8

–28.1

–27.7

n-C

–27.7 –28.6 –28.0 –27.2 –26.6 –28.1 –27.0 –28.6 –28.8 –28.7 –28.1

n-C

–27.7 –28.7 –27.8 –27.4 –26.5 –28.4 –26.9 –28.2 –27.9 –28.0 –27.8

n-C

–28.1 –28.1 –28.6 –27.6 –27.0 –29.2 –27.1 –28.8 –29.2 –28.8 –28.8

n-C

–28.3 –27.6 –29.1 –28.2 –26.7 –28.7 –27.0 –28.6 –28.0 –28.4 –27.9

n-C

–28.8

–28.9

–28.6

–28.1

–27.5

–29.1

–27.1

–28.9

–29.3

–28.9

n-C

–29.5 –28.4 –28.5 –27.8 –27.7 –29.1 –26.9 –28.7 –28.5 –28.3 –28.1

n-C

–29.7 –28.7 –28.5 –28.4 –28.2 –29.3 –26.9 –29.1 –29.4 –29.2 –28.9

n-C

–29.1

–26.4 –28.5 –27.8 –28.2 –28.5

n-C

–29.2 –27.2

–27.6

–27.6 –27.8

Table 3: Compound-specific carbon isotopic composition of selected samples (

C, ‰; PDB).

which argues for deposition of organic matter in a strongly
oxygen depleted environment.

Steroids and hopanoids

All the studied sediments are characterized by the occur-

rence of C

—C

regular steranes in very low concentrations.

For  the  most  samples,  the  steroid  contents  vary  between  0.1
and  0.6 µg/g TOC  (Table 2),  although  higher  concentrations
(0.8—1.3 µg/g TOC)  were  also  found,  mostly  in  the  middle
and upper parts of the core samples from unit III. Only sample
C-14 is characterized by raised steroid contents (up to 5.6 µg/g
TOC;  Table 2).  Interestingly,  diasteranes  or  diasterenes  were
not detected, despite the high amounts of clay in the samples.
Furthermore,  the  aromatic  fraction  is  characterized  by  com-
plete absence of aromatic steroids indicating the low maturity
of the organic matter (Mackenzie et al. 1981).

In all samples, comparable steroid distribution patterns,

composed of either nearly equal amounts of C

and C

steranes, or a slightly predominating C

compounds, were de-

tected (Table 2, Fig. 8). For most samples, the C

steranes

are present in very low amounts (10—20 %; Table 2). Slightly
increased C

concentrations (up to 40 %) exist only within

the core samples from unit III. The steranes show very low
20S/(20S + 20R) ratios of the

ααα isomers (Fig. 8). In addi-

tion to the regular steranes, a series of C

—C

4-methylster-

anes in almost equal amounts were detected in sample C-14.

Steranes originate from sterols present in living organisms

(Mackenzie et al. 1982; Waples & Machihara 1991; Meyers
1997; Volkman 2003, 2005). The main sources of C

sterols

are photosynthesizing organisms, including terrestrial plants,
while C

sterols are more characteristic for marine phy-

toplankton (Brassell & Eglinton 1981; Volkman 1986;
Barrett et al. 1995; Volkman et al. 1998). However, numer-
ous results from biomarker studies indicate that phytoplank-
ton and microalgae also produce high proportions of C

sterols (Volkman et al. 1998; Volkman 1999, 2003). The

421

ORGANIC MATTER AND HYDROCARBON IN MIDDLE MIOCENE SEDIMENTS (NE BULGARIA)

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

Fig. 8. Typical mass chromatogram (m/z 217 + 231) of the steroids and their methyl substituted derivatives.

most probable source of C

steranes are marine microalgae,

as suggested by the presence of significant amounts of C

sterols in marine diatoms (Kates et al. 1977; Barrett et al.
1995; Volkman et al. 1998; Volkman 2003). Methyl steroids
are also related to marine and lacustrine microalgal precursors
(Volkman et al. 1990, 1998). A specific class of C

4-me-

thyl steranes with dinosterol structure is a biomarker for dino-
flagellates (Robinson et al. 1984; Mansour et al. 1999).

Considering these facts, the steranes distribution patterns ar-

gue for mixed organic matter, composed predominantly of
marine phytoplankton, but with significant contributions from
terrestrial plants. Enhanced contributions of C

steranes are

observed  in  few  core  samples  from  unit III  (e.g.  C-1,  C-3,
C-14;  Table 2),  as  well  as  in  sample  C-31  from  unit IIa,
(Table 2). None of these samples shows increased concentra-
tion  of  long-chain  n-alkanes.  Therefore,  the  slight  predomi-
nance of C

steranes in these samples might reflect a change

in the phytoplankton communities. It is noticeable that in con-
trast to n-alkanes distribution, the steranes distributions are
similar in outcrop and core samples from unit III. Therefore,
significant compositional changes in the short-chain hydrocar-
bons, possibly due to weathering, can be assumed.

Hopane concentrations range from 1.5 to 11.8 µg/g TOC

(Table 2, Fig. 4). Two set of samples can be separated based
on the hopane composition. The first set comprises the core
samples from unit III, in which hopanes are represented by
17

α,21β(H)-, 17β,21α(H), and 17β,21β(H) compounds with

27—32 carbon atoms (Fig. 9a). C

hopanes are missing. An

overall  positive  correlation  in  these  samples  exists  between
the hopane and sterane concentrations (r = 0.76), suggesting a
proportional  increase  of  bacterial  activity  together  with  the
increase of organic matter input (see also Fig. 4). The second
set  comprises  samples  from  subunit IIa  and  the  outcrop
samples, in which hopanes are represented by 17

α,21β(H)-,

β,21α(H) and 17β,21β(H) compounds with 27—35 carbon

atoms, with the C

hopane absent (Fig. 9b). For subunit IIa

the vertical distribution of the hopanoids shows a marked
upward increase (Fig. 4). However, no correlation exists be-

tween the hopane and sterane concentrations, suggesting that
the increase of bacterial activity is probably related to a shift
to a more oxygen-rich environment, as recent studies indi-
cated aerobic bacteria as the main source of hopanes (Talbot
et al. 2008; Rezanka et al. 2010).

For both sets the hopane distribution is characterized by a

marked predominance of hopanoids in the 17

α(H),21β(H)

configuration (Fig. 9 – Ourisson et al. 1979). Dominant in all
samples are 17

α(H)-C

and 17

α(H)-C

hopanes (Fig. 9), ac-

companied by 17

β,21α(H)—C

moretane and C

hopanes.

The core samples from unit III are characterized by low
Ts/(Ts+Tm) ratios (0.2—0.4; Table 2). This ratio and appreci-
able amounts of hopanes with 17

β,21β(H) configuration

(Fig. 9a) argue for a low maturity of the organic matter. In con-
trast, Ts/(Ts + Tm) ratio is significantly higher (0.4—0.8;
Table 2) for subunit IIa samples and the outcrop samples, in-
dicating enhanced diagenetic changes. Furthermore, the ratio
of 22S/(22S + 22R) C

-hopanes shows generally very high

values (0.38—0.66), which are typical for marginal mature
organic matter. However, both ratios are known to be influ-
enced by facies variations (Waples & Machihara 1991; Peters
et al. 2005).

The most probable biological precursors of hopanoid

biomarkers  are  bacteriohopanepolyol  derivatives,  and  to  a
lesser extent 3-desoxyhopanes (Ourisson et al. 1979; Rohmer
et  al.  1992),  identified  in  the  membrane  of  aerobic  bacteria
such  as  cyanobacteria,  and  heterotrophic  and  methano-
trophic bacteria (e.g. Gibson et al. 2008; Talbot et al. 2008;
Rezanka et al. 2010).

The extended hopanes with 31 to 35 carbon atoms are rep-

resented in relatively low amounts (data not shown) (Fig. 9).
According to Van Dorselaer et al. (1975), the formation of
17

α,21β(H)-homohopane (22R) implies complex reactions

in acidic environments under oxic conditions. However, a
direct bacterial input to the organic matter should also be
considered, as Thiel et al. (2003) has proved the presence of
homohopanoic acids with

αβ configuration in living micro-

bial mats.

422

ZDRAVKOV, BECHTEL, ĆORIĆ and SACHSENHOFER

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

Hop-17(21)-ene was detected in small amounts ( < 1.9 µg/g

TOC; Table 2) in almost all samples. Its biological precursor
has not yet been established (Bottari et al. 1972; Ourisson et
al. 1979; Brassell et al. 1980; Volkman et al. 1986; Wakeham
et al. 1980b; Wolff et al. 1992; Sabel et al. 2005).

Other terpenoids with non-hopanoid skeleton

Terpenoids are widespread in the geosphere (Thomas

1969) as part of the tissues of higher plants. Among them,

Fig. 9. Typical mass chromatograms (m/z 191) of the hopanoid hydrocarbons.

di-and  tri-terpenoids  are  valuable  biomarkers  for  determin-
ing  the  type  of  precursor  vegetation  (Simoneit  1977,  1986,
1999).  In  most  of  the  studied  samples  very  low  amounts
(<1.3 µg/g TOC;  Table 2)  of  diterpenoids  were  detected.
They  are  represented  by  hydrocarbons  with  phyllocladane
and  pimarane  type  structure.  The  individual  compounds
identified are

α-phyllocladane, occurring as a major compo-

nent, whereas the samples from subunit IIa and the outcrop
samples also contain trace amounts of

β-phyllocladane and

pimarane. Phyllocladanes are biomarkers for gymnosperms

423

ORGANIC MATTER AND HYDROCARBON IN MIDDLE MIOCENE SEDIMENTS (NE BULGARIA)

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

Fig. 10. a – Typical gas chromatogram of the aromatic fraction: MN – methyl naphthalenes, DMN – dimethyl naphthalenes, TMN – tri-
methyl naphthalenes, Ph – phenanthrene, MPh – methyl phenanthrenes; b – Gas chromatogram of the aryl isoprenoids (m/z 133) and chro-
mans (m/z 400 + 414): ar-i-C

14 ÷ 19

trimethyl (2,3,6)-alkylbenzol, Di-MTTC – dimethylated 2-methyl-2-(4’,8’,12’-trimethyltridecyl)chroman;

Tri—MTTC – trimethylated 2-methyl-2-(4’,8’,12’-trimethyltridecyl)chroman.

(e.g.  Podocarpaceae,  Taxodiaceae,  Araucariaceae  and
Cupressaceae – Noble et al. 1985; Sukh Dev 1989; ten Haven
et al. 1992; Otto & Wilde 2001). Triterpenoids with oleanane,
lupane  and  ursane  type  skeleton  structure  and  their  aromatic
derivatives,  derived  from  angiosperm  plants  (Philp  1985;
Simoneit 1986), were not detected in the studied samples.

Aromatic hydrocarbons

The composition of the aromatic fraction in all the studied

samples  is  relative  uniform  and  comprises  methyl-,  dim-
ethyl- and trimethyl-naphthalenes, phenanthrene and its me-
thyl  derivatives,  as  well  as  low  amounts  of  polyaromatic
hydrocarbons (PAHs), represented by fluoranthrene, pyrene,
benzo[a]pyrene and chrysene (Table 2, Fig. 10a). PAHs may
form  due  to  diagenetic  alteration  of  natural  biolipids,  but
most authors accept an origin by wildfires (e.g. LaFlamme &
Hites  1979;  Wakeham  et  al.  1980a,b;  Venkatesan  &  Dahl
1989; Killops & Massoud 1992; Yunker 2002, 2003). Con-
sidering this, we speculate that the detected PAHs are of ter-
restrial  origin  and  were  transported  into  the  basin  together
with charcoal particles.

The aromatic hydrocarbon composition is further charac-

terized by the occurrence of aryl isoprenoids in the range
C

—C

with a 2,3,6-trimethyl substitution pattern for the

aromatic  ring  and  a  tail-to-tail  isoprenoid  chain  (Fig. 10b).
They  are  present  in  low  amounts  in  the  core  samples  from
unit III  (Table 2).  The  aryl  isoprenoids  are  derived  from
isorenieratene, which is known to be synthesized by photo-
synthetic  green  sulphur  bacteria.  These  organisms  are  pho-
totrophic anaerobes and require both light and H

S for growth

(Summons 1993; Pfennig 1997). Therefore, aryl isoprenoids
are  widely  used  as  a  biomarker  for  green  sulphur  bacteria
and  photic  zone  anoxia  (e.g.  Summons  &  Powell  1987;
Sinninghe  Damsté  et  al.  2001).  However,  Koopmans  et  al.
(1996) found out that aryl isoprenoids with a 2,3,6-trimethyl
substitution pattern can also form due to diagenetic transfor-
mation  of  the  more  ubiquitous  carotenoid

β-carotene, and,

therefore, have only limited applicability as a biomarker. In
the present case, isorenieratene itself has not been identified
in the samples.

The aromatic fraction of the samples from unit III is fur-

ther characterized by very low amounts of di- and trimethyl-
ated 2-methyl-2-trimethyltridecylchromans (MTTC; Table 2,
Fig. 10b). Among them, tri-MTTCs predominate (di-/tri-
MTTCs ratios: 0.1—0.4; Table 2). There are, however, samples
(e.g. C-5, C-7, C-11), which also contain appreciable amounts
of di-MTTCs (di-/tri-MTTCs ratios: 0.6—0.7).

The origin of the methyl substituted MTTCs is still un-

known  (e.g.  Li  et  al.  1995).  Sinninghe  Damsté  et  al.  (1987,
1993) suggest that MTTCs form in the upper parts of the water
column, either from photosynthetic or non-photosynthetic or-
ganisms.  On  the  contrary,  Barakat  &  Rullkötter  (1997)  sug-
gested  that  chromans  are  a  result  of  cyclization  of  alkylated
phenols. Despite this limitation, methyl-substituted chromans
are widely used as a paleosalinity indicators (e.g. Schwark &
Püttmann 1990; Sachsenhofer et al. 2009). Sinninghe Damsté
et al. (1987, 1993) indicate that formation and preservation of
mono-, di- and trimethyl substituted chromans require marine
environments with normal to enhanced salinity and stratifica-
tion  of  the  water  column.  Furthermore,  the  authors  observed
increasing  di-/tri-MTTC  ratios  with  increasing  salinity.  Con-

424

ZDRAVKOV, BECHTEL, ĆORIĆ and SACHSENHOFER

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

sidering this, fully marine conditions are suggested during the
deposition of unit III, perhaps with a salinity stratified water
column. This suggestion is supported by the lamination of the
clays, as well as by the overall positive correlation between
the concentrations of MTTC and sulphur (r = 0.51) as well as
MTTC and aryl isoprenoids (r = 0.68).

Conclusions

The Middle to Upper Miocene succession in the Varna-

Balchik Depression, located in the south-eastern part of the
Moesian Platform, is up to 300 m thick and is represented by
sandy, clayey and carbonate rocks. Based on lithology and
resistivity log response, the succession is subdivided from
base to top into five units.

Sedimentation of fine-grained siliciclastic material pre-

vailed  during  deposition  of  the  lower  parts  of  units I  and  II
(subunits Ia and IIa, respectively), whereas their upper parts
(subunits Ib,  IIb)  are  composed  of  biomicritic  limestone.
Based on high diatom contents, subunit IIa is considered co-
eval  with  the  lower  part  of  the  Euxinograd  Formation.
Unit III represents the upper part of the Euxinograd Forma-
tion. It is composed of laminated clays with bioclastic layers
and significant amounts of biogenic silica, which are mainly
derived from sponge spicules in core samples and from dia-
tom frustules in outcrop samples. Units IV and V are repre-
sented predominantly by carbonate rocks and are correlated
with the Upper Miocene Topola and Karvuna Formations.

Whereas TOC contents in subunit Ia are negligible

(<0.2 wt. %),  TOC  contents  in  subunit IIa  and  unit III
(Euxinograd  Fm)  are  generally  between  1  and  2 wt. %,  but
can reach 4 wt. % in unit III. Sulphur contents are typically
low  ( < 0.5 %)  and  exceed  1.0 %  only  in  samples  from
unit III.  Low  sulphur  contents  may  be  due  to  deposition  in
environments  with  reduced  salinity.  Normal  marine  condi-
tions are suggested for unit III.

The biomarker composition is typical for mixed marine and

terrestrial  organic matter, and shows no significant  changes
in  subunit IIa  and  unit III.  The  molecular  composition  and
biomarker ratios support the presence of immature organic mat-
ter,  deposited  in  dysoxic  to  anoxic  environments.  Kerogen  is
mainly type II in subunit IIa (average HI = 324 mg HC/g TOC)
and type III for unit III (average HI ~ 200 mg HC/g TOC).

Considering the classification scheme of Peters (1986),

TOC and Rock Eval data show that subunit IIa holds a fair to
good hydrocarbon potential for oil, whereas the upper 5 m of
unit III holds a good to fair potential with the possibility to
generate gas and minor oil. Both units are immature in on-
shore areas, but thickness and maturity may increase towards
the east in offshore areas.

Acknowledgments: Financial support from OMV, awarded
to AZ, is greatly appreciated. In addition, the authors would
like  to  express  their  gratitude  to  E.  Kozhukharov,  OMV
Bulgaria,  for  his  technical  assistance.  The  manuscript  also
benefited  from  the  valuable  suggestions  made  by  Dr.  P.
Kosakowski, Dr. M. Stefanova, Dr. E. Koleva-Rekalova and
an anonymous reviewer.

References

Albro P.W. & Dittmer J.C. 1967: Bacterial hydrocarbons: occurrence,

structure and metabolism. Lipids 5, 320—325.

Allen J.E., Formery F.W. & Markovetz A.J. 1971: Microbial degrada-

tion of n-alkanes. Lipids 6, 448—452.

Barakat A.O. & Rullkötter J. 1997: A comparative study of molecular

paleosalinity indicators: chromans, tocopherols and C20 iso-
prenoid thiophenes in Miocene lake sediments (Nördlinger Ries,
Southern Germany). Aquat. Geochem. 3, 169—190.

Barrett S., Volkman J.K. & Dunstan G.A. 1995: Sterols of 14 Species

of marine diatoms (bacillariophyta). J. Phycology 31, 360—369.

Berner R.A. 1984: Sedimentary pyrite formation: An update. Geochim.

Cosmochim. Acta 48, 605—615.

Blumer M., Guillard R.R. & Chase T. 1971: Hydrocarbons of marine

phytoplankton. Int. J. Life Ocean. Coast. Waters 8, 180—183.

Bottari F., Marsili A., Morelli I. & Pacchiani M. 1972: Aliphatic and tri-

terpenoid hydrocarbons from ferns. Phytochemistry 11, 2519—2523.

Brassell S.C. & Eglinton G. 1981: Biogeochemical significance of a

novel sedimentary C27 stanol. Nature 290, 579—582.

Brassell S.C., Comet P.A., Eglinton G., Isaacson P.J., McEvoy J., Max-

well J.R., Thompson I.D., Tibbetts P.J.C. & Volkman J.K. 1980:
The origin and fate of lipids in the Japan Trench. In: Douglas A.G.
& Maxwell J.R. (Eds.): Advances in organic geochemistry 1979.
Pergamon Press, Oxford, 375—392.

Bray E.E. & Evans E.D. 1961: Distribution of n-paraffins as a clue to

recognition of source beds. Geochim. Cosmochim. Acta 22, 2—15.

Brooks J.D., Gould K. & Smith J.W. 1969: Isoprenoid hydrocarbons in

coal and petroleum. Nature 222, 257—259.

Cheshitev G., Chontova T., Popov N. & Kojumdgieva E. 1995: Explana-

tory note to the geological map of Bulgaria on scale 1 : 100,000,
Balchik and Shabla map sheets. “Geology and Geophysics” AG
(in Bulgarian).

Cheshitev G., Chontova T., Milanova V., Popov N. & Kojumdgieva E.

1994a: Explanatory note to the geological map of Bulgaria on
scale 1 : 100,000, Dobrich map sheet. “Geology and Geophysics” AG
(in Bulgarian).

Cheshitev G., Milanova V., Popov N. & Kojumdgieva E. 1994b: Ex-

planatory note to the geological map of Bulgaria on scale
1 : 100,000, Varna and Zlatni Pyassatsi map sheets. “Geology and
Geophysics” AG (in Bulgarian).

Cranwell P.A. 1977: Organic geochemistry of CamLoch (Sutherland)

sediments. Chem. Geol. 20, 205—221.

Dabovski C., Boyanov I., Khrischev K., Nikolov T., Sapunov I., Yanev

Y. & Zagorchev I. 2002: Structure and Alpine evolution of Bul-
garia. Geol. Balcanica 32, 9—15.

Darakchieva S. 1989: Foraminiferal zonation of the Miocene in North-

eastern Bulgaria. Bulg. Acad. Sci. Palaeontol. Stratigr. Lithol. 27,
31—43 (in Bulgarian).

DeMaster D. 2003: The diagenesis of biogenic silica: Chemical trans-

formations occurring in the water column, seabed, and crust. In:
Mackenzie F.T. (Ed.): Sediments, diagenesis, and sedimentary rocks.
Treatise on Geochemistry 7, Elsevier Pergamon, 87—98.

Didyk B.M., Simoneit B.R.T., Brassell S.C. & Eglinton G. 1978: Or-

ganic geochemical indicators of paleoenvironmental conditions of
sedimentation. Nature 272, 216—222.

Eglinton G. & Hamilton R.J. 1967: Leaf epicuticular waxes. Science 156,

1322—1335.

Espitalié J., Marquis F. & Barsony I. 1984: Geochemical logging. In:

Voorhess K.J. (Ed.): Analytical pyrolysis. Butterworths, Boston,
276—304.

Espitalié J., LaPorte J.L., Madec M., Marquis F., Leplat P., Paulet J. &

Boutefeu A. 1977: Méthode rapide de la caractérisation des roches
méres de leur potential pétrolier et de leur degré dévolution. Rev.
Inst. France Pétrole 32, 23—42.

Ficken K.J. 2000: An n-alkane proxy for the sedimentary input of sub-

merged/floating freshwater aquatic macrophytes. Org. Geochem.
31, 745—749.

Gagosian R., Nigrelli G. & Volkman J.K. 1983: Vertical transport and

425

ORGANIC MATTER AND HYDROCARBON IN MIDDLE MIOCENE SEDIMENTS (NE BULGARIA)

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

transformation of biogenic organic compounds from a sediment
trap experiment off the coast of Peru. In: Coastal upwelling and its
sediment record. NATO Conf. Series 10B, 241—272.

Georgiev G. 2012: Geology and hydrocarbon systems in the Western

Black Sea. Turkish J. Earth Sci. 21, 723—754.

Gibson R., Talbot H., Kaur G., Pancost R. & Mountain B. 2008:

Bacteriohopanepolyol signatures of cyanobacterial and methano-
trophic bacterial populations recorded in a geothermal vent sinter.
Org. Geochem. 39, 1020—1023.

Goossens H., de Leeuw J.W., Schenck P.A. & Brassell S.C. 1984: To-

copherols as likely precursors of pristane in ancient sediments and
crude oils. Nature 312, 440—442.

Han J. & Calvin M. 1969: Hydrocarbon distribution of algae and bacte-

ria, and microbiological activity in sediments. Proc. Nat. Acad.
Sci. U.S.A. 64, 436—443.

Han J., McCarthy E.D., Van Hoeven W., Calvin M. & Bradley W.H.

1968: Organic geochemical studies. II. A preliminary report on
the distribution of aliphatic hydrocarbons in algae, in bacteria, and
in a recent lake sediment. Proc. Nat. Acad. Sci. U.S.A. 29, 29—33.

Hughes W.B., Holba A.G. & Dzou L.I.P. 1995: The ratios of dibenzo-

thiophene to phenanthrene and pristane to phytane as indicators of
depositional environment and lithology of petroleum source rocks.
Geochim. Cosmochim. Acta 59, 3581—3598.

Johnson R.W. & Calder J.A. 1973: Early diagenesis of fatty acids and

hydrocarbons in a salt marsh environment. Geochim. Cosmochim.
Acta 37, 1943—1955.

Kates M., Tremblay P., Volcani B.E., Diego S. & Jolla L. 1977: Identifi-

cation of the free and conjugated sterol in a non-photosynthetic dia-
tom, Nitzschio alba, as 24-methylene cholesterol. Lipids 13, 34—41.

Killops S. & Massoud M. 1992: Polycyclic aromatic hydrocarbons of

pyrolytic origin in ancient sediments: evidence for Jurassic vege-
tation fires. Org. Geochem. 18, 1—7.

Kojumdgieva E. 1965: On some properties of the stratigraphy of

Karaganian age in Varna district. Bull. Inst. Sci. Rech. Geol. 2,
49—59 (in Bulgarian).

Kojumdgieva E. & Dikova P. 1978: Miocene sediments from borings

near Obrochishte (North—Eastern Bulgaria). Bulg. Acad. Sci.
Palaeontol. Stratigr. Lithol. 8, 69—75 (in Bulgarian).

Kojumdgieva E. & Popov N. 1981: Regions structuro-paleogeogra-

phiques en Bulgarie du Nord—Est pendant le Negene. C.R. Acad.
Bulg. Sci. 34, 1273—1275.

Kojumdgieva E. & Popov N. 1987. The molluscan communities from

the Sarmatian in North—Eastern Bulgaria. Bulg. Acad. Sci.
Palaeontol. Stratigr. Lithol. 24, 60—69 (in Bulgarian).

Koleva-Rekalova E. 1994: Sarmatian aragonite sediments in Northeast-

ern Bulgaria – origin and diagenesis. Geol. Balcanica 24, 47—64.

Koleva-Rekalova E. 1997: Sedimentological characteristic of the Sar-

matian rocks from the Balchik area, NE Bulgaria. Rev. Bulg. Geol.
Soc. 58, 31—42 (in Bulgarian).

Koleva-Rekalova E. 1998: Formational conditions of the Sarmatian

sedimentary rocks from the Balchik area, North—Eastern Bulgaria.
Rev. Bulg. Geol. Soc. 59, 1, 69—74 (in Bulgarian).

Koleva-Rekalova E. & Darakchieva S. 2002: The Bessarabian foramin-

iferal limestones of the Odurtsi Formation, Northeastern Bulgaria.
Geol. Miner. Resour. 4, 31—33 (in Bulgarian).

Koopmans M.P., Schouten S., Kohnen M.E.L. & Sinninghe Damsté

J.S. 1996: Restricted utility of aryl isoprenoids as indicators for
photic zone anoxia. Geochim. Cosmochim. Acta 60, 4873—4876.

Koopmans M.P., Rijpstra W.I.C., Klapwijk M.M., de Leeuw J.W.,

Lewan M.D. & Sinninghe Damsté J.S. 1999: A thermal and chemi-
cal degradation approach to decipher pristane and phytane precur-
sors in sedimentary organic matter. Org. Geochem. 30, 1089—1104.

Krautter M. 1998: Ecology of siliceous sponges – application to the

environmental interpretation of the Upper Jurassic sponge facies
(Oxfordian) from Spain. Cuad. Geol. Ibérica 24, 223—239.

LaFlamme R.E. & Hites R.A. 1979: Tetra- and pentacyclic, naturally

occurring, aromatic hydrocarbons in recent sediments. Geochim.
Cosmochim. Acta 43, 1687—1691.

Langford F.F. & Blanc-Valleron M.-M. 1990: Interpreting Rock-Eval

pyrolysis data using graphs of pyrolyzable hydrocarbons vs. total

organic carbon. AAPG Bulletin 74, 799—804.

Li M., Larter S., Taylor P., Jones D., Bowler B. & Bjoroy M. 1995:

Biomarkers or not biomarkers? A new hypothesis for the origin of
pristane involving derivation from methyltrimethyltridecyl-
chromans (MTTCs) formed during diagenesis from chlorophyll
and alkylphenols. Org. Geochem. 23, 159—167.

Lichtfouse E., Derenne S., Mariotti A. & Largeau C. 1994: Possible al-

gal origin of long chain odd n-alkanes in immature sediments as re-
vealed by distributions and carbon isotope ratios. Org. Geochem.
22, 1023—1027.

Lourens L., Hilgen F., Shackleton N.J., Laskar J. & Wilson J. 2004:

The Neogene period. In: Gradstein F., Ogg J. & Smith A. (Eds.):
A Geologic Time Scale 2004. Cambridge University Press, Cam-
bridge, 409—440.

Mackenzie A.S., Hoffmann C.F. & Maxwell J.R. 1981: Molecular para-

meters of maturation in the Toarcian shales, Paris Basin, France. III.
Changes in aromatic steroid hydrocarbons. Geochim. Cosmochim.
Acta 45, 1345—1355.

Mackenzie A.S., Brassell S.C., Eglington G. & Maxwell J.R. 1982: Chemi-

cal fossils: the geological fate of steroids. Science 80, 217, 491—504.

Mansour M., Volkman J.K., Jackson A.E. & Blackburn S.I. 1999: The

fatty acid and sterol composition of five marine dinoflagellates. J.
Phycology 35, 710—720.

Meyers P. 1994: Preservation of elemental and isotopic source identifi-

cation of sedimentary organic matter. Chem. Geol. 114, 289—302.

Meyers P. 1997: Organic geochemical proxies of paleoceanographic,

paleolimnologic, and paleoclimatic processes. Org. Geochem. 27,
213—250.

Milakovska Z., Koleva-Rekalova E. & Rekalov K. 2001: A new occur-

rence of tephra layers in the Lower Sarmatian (Volhynian) clay
deposits from NE Bulgaria. Rev. Bulg. Geol. Soc. 62, 31—35.

Noble R.A., Alexander R., Kagi R.I. & Knox J. 1985: Tetracyclic

diterpenoid hydrocarbons in some Australian coals, sediments and
crude oils. Geochim. Cosmochim. Acta 49, 2141—2147.

Ori G.G. 2004: The Miocene depositional setting of northeastern Bul-

garia. Unpubl. Report, 1—75.

Otto A. & Wilde V. 2001: Sesqui-, di-, and triterpenoids as chemosystem-

atic markers in Extant conifers – a review. Bot. Rev. 67, 141—238.

Ourisson G., Albrecht P. & Rohmer M. 1979: The hopanoids: Palaeo-

chemistry and biochemistry of a group of natural products. Pure
Appl. Chem. 51, 709—729.

Peters K.E. 1986: Guidelines for evaluating petroleum source rock us-

ing programmed pyrolysis. AAPG Bulletin 70, 318—329.

Peters K.E., Walters C.C. & Moldowan J.M. 2005: The biomarker guide

Vol. 2: Biomarkers and isotopes in petroleum exploration and Earth
history. Cambridge University Press, New York, 1—1155.

Pfennig N. 1997: Phototrophic green and purple bacteria: a compara-

tive, systematic survey. Ann. Rev. Microbiol. 31, 275—290.

Philp R.P. 1985: Fossil fuel biomarkers. Applications and spectra.

Methods Geochem. Geophys. 23, 1—294.

Popov N. & Kojumdgieva E. 1987: The Miocene in Northeastern Bul-

garia (lithostratigraphic subdivision and geological evolution).
Rev. Bulg. Geol. Soc. 38, 15—33 (in Bulgarian).

Popov N., Stancheva M. & Darakchieva S. 1986: Reference sections of

the Neogene from Northeastern Bulgaria. Bulg. Acad. Sci. Palae-
ontol. Stratigr. Lithol. 23, 25—45 (in Bulgarian).

Radke M., Willsch H. & Welte D.H. 1980: Preparative hydrocarbon

group type determination by automated medium pressure liquid
chromatography. Anal. Chem. 52, 406—411.

Raiswell R. & Berner R.A. 1986: Pyrite and organic-matter in Phanerozoic

normal marine shales. Geochim. Cosmochim. Acta 50, 1967—1976.

Rezanka T., Siristova L., Melzoch K. & Sigler K. 2010: Hopanoids in bac-

teria and cyanobacteria – their role in cellular biochemistry and phy-
siology, analysis and occurrence. Mini-Rev. Org. Chem. 7, 300—313.

Ricken W. 1991: Variation of sedimentation rates in rhythmically bed-

ded sediments. Distinction between depositinal types. In: Einsele
G., Ricken W. & Seilacher A. (Eds.): Cycles and events in strati-
graphy. Springer, Berlin, 167—187.

Robinson N., Eglinton G., Brassell S.C. & Cranwell P.A. 1984:

Dinoflagellate origin for sedimentary 4a-methylsteroids and 5b

426

ZDRAVKOV, BECHTEL, ĆORIĆ and SACHSENHOFER

GEOL

EOL

EOLOGICA CARPA

OGICA CARPA

OGICA CARPATHICA

THICA

THICA, 2015, 66, 5, 409—426

(H)-stanols. Nature 308, 439—442.

Rohmer M., Bisseret P. & Neunlist S., 1992: The hopanoids, prokaryo-

tic triterpenoids and precursors of ubiquitous molecular fossils. In:
Moldowan J.M., Albrecht P. & Philp R.P. (Eds.): Biological
markers in sediments and petroleum. Prentice—Hall, Englewood
Cliffs, NJ, 1—17.

Sabel M., Bechtel A., Püttmann W. & Hoernes S. 2005: Palaeo-

environment of the Eocene Eckfeld Maar lake (Germany): impli-
cations from geochemical analysis of the oil shale sequence. Org.
Geochem. 36, 873—891.

Sachsenhofer R.F., Stummer B., Georgiev G., Dellmour R., Bechtel A.,

Gratzer R. & Ćorić S. 2009: Depositional environment and hydrocar-
bon source potential of the Oligocene Ruslar Formation (Kamchia
Depression; Western Black Sea). Mar. Petrol. Geol. 26, 57—84.

Schwark L. & Püttmann W. 1990: Aromatic hydrocarbon composition

of the Permian Kupferschiefer in the Lower Rhine Basin, N.W.
Germany. Org. Geochem. 16, 749—761.

Simoneit B.R.T. 1977: Diterpenoid compounds and other lipids in deep-

sea sediments and their geochemical significance. Geochim. Cos-
mochim. Acta 41, 463—476.

Simoneit B.R.T. 1986: Cyclic terpenoids of the geosphere. In: Johns

R.B. (Ed.): Biological markers in the sedimentary record. Elsevier,
Amsterdam, 43—99.

Simoneit B.R.T. 1999: A review of biomarker compounds as source in-

dicators and tracers for air pollution. Environ. Sci. Pollut. Res. 6,
159—169.

Sinninghe Damsté J.S., Schouten S. & van Duin A.C.T. 2001:

Isorenieratene derivatives in sediments: Possible controls on their
distribution. Geochim. Cosmochim. Acta 65, 1557—1571.

Sinninghe Damsté J.S., Kock-Van Dalen A., de Leeuw J.W., Schenck

P., Guoying S. & Brassell S.C. 1987: The identification of mono-,
di- and trimethyl 2-methyl-2-(4,8,12-trimethyltridecyl)chromans
and their occurrence in the geosphere. Geochim. Cosmochim. Acta
51, 2393—2400.

Sinninghe Damsté J.S., Keely B.J., Betts S.E., Baas M., Maxwell J.R.

& de Leeuw J.W. 1993: Variations in abundances and distribu-
tions of isoprenoid chromans and long-chain alkylbenzenes in
sediments of the Mulhouse Basin: a molecular sedimentary record
of palaeosalinity. Org. Geochem. 20, 1201—1215.

Steininger F.F. & Wessely G. 2000: From the Tethyan Ocean to the

Paratethys  Sea:  Oligocene  to  Neogene  stratigraphy,  paleogeo-
graphy  and  palaeobiogeography  of  the  circum-Mediterranean  re-
gion  and  the  Oligocene  to  Neogene  Basin  evolution  in  Austria.
Mitt. Österr. Geol. Gesell. 92, 95—116.

Sukh Dev 1989: Terpenoids. In: Rowe J.W. (Ed.): Natural products of

woody plants. Springer, Berlin, 691—807.

Summons R.E. 1993: Biogeochemical cycles: a review of fundamental

aspects of organic matter formation, preservation, and composi-
tion. In: Engel M.H. & Macko S.A. (Eds.): Organic geochemistry
– principles and applications. Plenum Press, NewYork, 3—21.

Summons R.E. & Powell T.G. 1987: Identification of aryl isoprenoids

in source rocks and crude oils: Biological markers for the green
sulphur bacteria. Geochim. Cosmochim. Acta 51, 557—566.

Talbot H., Summons R., Jahnke L., Cockell C., Rohmer M. &

Farrimond P. 2008: Cyanobacterial bacteriohopanepolyol signa-
tures from cultures and natural environmental settings. Org.
Geochem. 39, 232—263.

Temniskova-Topalova D. 1990: Biostratigraphic subdivision of Mio-

cene in northeast Bulgaria based on diatomic algae. In: Microfos-
sils in Bulgarian stratigraphy. Union of Researchers in Bulgaria,
Sofia, 73—80 (in Bulgarian).

ten Haven L., Peakman M. & Rullkötter J. 1992: Early diagenetic

transformation of higher-plant triterpenoids in deep-sea sediments
from Baffin Bay. Geochim. Cosmochim. Acta 56, 2001—2024.

ten Haven L., de Leeuw W.J., Rullkötter J. & Sinnighe Damsté J. 1987:

Restricted utility of the pristane/phytane ratio as a palaeoenviron-
mental indicator. Nature 330, 641—643.

Thiel V., Blumenberg M., Pape T., Seifert R. & Michaelis W. 2003:

Unexpected occurrence of hopanoids at gas seeps in the Black
Sea. Org. Geochem. 34, 81—87.

Thomas B.R. 1969: Kauri resins – modern and fossil. In: Eglinton G.

& Murphy M.T.J. (Eds.): Organic geochemistry – methods and
results. Springer, New York, 599—618.

Toledo F.A.L., Cachão M., Costa K.B. & Pivel M.A.G. 2007: Plank-

tonic foraminifera, calcareous nannoplankton and ascidian varia-
tions during the last 25 kyr in the Southwestern Atlantic: A
paleoproductivity signature? Mar. Micropaleont. 64, 67—79.

Tucker M.E. & Wright V.P. 1990: Carbonate sedimentology. Blackwell

Science Ltd., 1—482.

Van Dorselaer A., Albrecht P. & Connan J. 1975: Changes in composi-

tion of polycyclic alkanes by thermal maturation (Yallourn Lig-
nite, Australia). In: Campus R. & Goni J. (Eds.): Advances in
organic geochemistry. Enadimsa, Madrid, 53—59.

Venkatesan M.I. & Dahl J. 1989: Organic geochemical evidence for glo-

bal fires at the Cretaceous/Tertiary boundary. Nature 338, 57—60.

Volkman J.K. 1986: A review of sterol markers for marine and terrig-

enous organic matter. Org. Geochem. 9, 83—99.

Volkman J.K. 1988: Biological marker compounds as indicators of the

depositional environments of petroleum source rocks. In: Fleet
A.J., Kelts K. & Talbot M.R. (Eds.): Lacustrine petroleum source
rocks. Geol. Soc. London, Spec. Publ. 40, 103—122.

Volkman J.K. 1999: Eustigmatophyte microalgae are potential sources

of C29 sterols, C22—C28 n-alcohols and C28—C32 n-alkyl diols in
freshwater environments. Org. Geochem. 30, 307—318.

Volkman J.K. 2003: Sterols in microorganisms. Appl. Microbiol. Bio-

technol. 60, 495—506.

Volkman J.K. 2005: Sterols and other triterpenoids: source specificity and

evolution of biosynthetic pathways. Org. Geochem. 36, 139—159.

Volkman J.K. & Maxwell J.R. 1986: Acyclic isoprenoids as biological

markers. In: Johns R.B. (Ed.): Biological markers in the sedimen-
tary record. Elsevier, Amsterdam, 1—42.

Volkman J.K. & Tanoue E. 2002: Chemical and biological studies of par-

ticulate organic matter in the ocean. J. Oceanography 58, 265—279.

Volkman J.K., Johns R. & Gillan F. 1980: Microbial lipids of an inter-

tidal sediment – I. Fatty acids and hydrocarbons. Geochim.
Cosmochim. Acta 44, 1133—1143.

Volkman J.K., Kearney P. & Jeffrey S.W. 1990: A new source of 4-me-

thyl sterols and 5 [alpha](H)-stanols in sediments: prymnesiophyte
microalgae of the genus Pavlova. Org. Geochem. 15, 489—497.

Volkman J.K., Farrington J., Gagosian R.B. & Wakeham S.G. 1983:

Lipid composition of coastal marine sediments from the Peru up-
welling region. In: Bjoroy M. (Ed.): Advances in organic geochem-
istry 1981. John Wiley and Sons, London, 228—240.

Volkman J.K., Allen D., Stevenson P. & Burton H. 1986: Bacterial and

algal hydrocarbons in sediments from a saline Antarctic lake, Ace
Lake. Org. Geochem. 10, 671—681.

Volkman J.K., Barrett S.M., Blackburn S.I., Mansour M.P., Sikes E.L.

& Gelin F. 1998: Microalgal biomarkers: A review of recent re-
search developments. Org. Geochem. 29, 1163—1179.

Wakeham S.G., Schaffner C. & Giger W. 1980a: Polycyclic aromatic

hydrocarbons in recent lake sediments. I. Compounds having an-
thropogenic origins. Geochim. Cosmochim. Acta 44, 403—413.

Wakeham S.G., Schaffner C. & Giger W. 1980b: Polycyclic aromatic

hydrocarbons in recent lake sediments – II. Compounds derived
from biological precursors during early diagenesis. Geochim.
Cosmochim. Acta 44, 415—429.

Waples D. & Machihara T. 1991: Biomarkers for geologists – a prac-

tical guide to the application of steranes and triterpanes in petro-
leum geology. Methods in Exploration, AAPG, 1—91.

Wolff G.A., Rukin N. & Marshall J.D. 1992: Geochemistry of an early

diagenetic concretion from the Birchi Bed (L. Lias, W. Dorset,
UK). Org. Geochem. 19, 431—444.

Yunker M. 2002: PAHs in the Fraser River basin: a critical appraisal of

PAH ratios as indicators of PAH source and composition. Org.
Geochem. 33, 489—515.

Yunker M. 2003: Alkane and PAH depositional history, sources and

fluxes in sediments from the Fraser River Basin and Strait of
Georgia, Canada. Org. Geochem. 34, 1429—1454.

Zolitschka B. 1998: Paläoklimatische Bedeutung laminierter Sedimente.

Relief Boden Paläoklima 13, 1—176.