Content uploaded by Jack J Middelburg
Author content
All content in this area was uploaded by Jack J Middelburg
Content may be subject to copyright.
INTRODUCTION
In marine sediments pyrite occurs as micron-sized euhedral crystals and
micron-sized crystals in a raspberry-like morphology (framboidal pyrite)
(e.g., Sweeney and Kaplan, 1973). Framboidal textures are the result of rapid
pyrite formation from aqueous solutions highly supersaturated with both Fe
monosulfides and pyrite, in which reaction kinetics favor the formation of Fe
monosulfides before pyrite. By contrast, euhedral pyrite forms more slowly at
saturation levels that are below those of Fe monosulfides (e.g., Sweeney and
Kaplan, 1973; Goldhaber and Kaplan, 1974; Raiswell, 1982; Rickard, 1997),
as confirmed by pyrite synthesis experiments (Wang and Morse, 1996) which
show that pyrite morphology changes from cubic to octahedral to spherulitic
with increasing degree of supersaturation. Raiswell (1982) argued that the tex-
ture of pyrite in pyritiferous carbonate concretions in ancient sedimentary
rocks depends on the relative rates of HS–generation and Fe supply. Bisulfide
generation is determined by the rate of SO42– reduction, and Fe supply by the
in situ availability of reactive iron. In these concretions, framboidal pyrite ap-
parently formed during early stages of diagenesis. During later stages of dia-
genesis, euhedral pyrite with higher 34S/32S isotope signatures formed because
in situ Fe sources and pore-water SO42– became depleted, and the SO42– re-
duction rate decreased. Wilkin et al. (1996) proposed that size distributions of
framboids are indicative of formation conditions: framboids found in sedi-
ments underlying euxinic water columns are smaller than those in sediments
underlying oxic or dysoxic water columns.
In this study, we assess the relation between the content, microtexture,
and sulfur isotope composition of pyrite within and below the most recent
eastern Mediterranean sapropel S1. We integrate the well-characterized pyrite
formation conditions known from laboratory studies (Sweeney and Kaplan,
1973; Rickard, 1997; Wang and Morse, 1996), the inferred formation path-
ways of pyrite in sediments and sedimentary rocks (Raiswell, 1982; Wilkin et
al., 1996), and recent insights into sulfur isotope fractionation during the ox-
idative part of the sulfur cycle (Canfield and Thamdrup, 1994). These com-
plementary approaches provide a new insight into the conditions of pyrite for-
mation and related redox conditions in eastern Mediterranean sediments
during sapropel formation.
EASTERN MEDITERRANEAN SEDIMENTS
Sediments in the eastern Mediterranean are characterized by the pres-
ence of organic-rich layers, sapropels, in a hemipelagic sediment sequence.
The origin of sapropels is still a matter of debate. Climate-induced enhanced
productivity (e.g., Calvert, 1983) and preservation due to oxygen depletion in
the bottom waters (e.g., Rossignol-Strick et al., 1982) are thought to be in-
volved in sapropel formation. The youngest sapropel, S1, formed between 5
and 9 ka (e.g., van Santvoort et al., 1996).
Sapropels are enriched in Fe and S as a result of bacterial SO42– reduc-
tion within the sapropels and pyrite formation during and shortly after sapro-
pel deposition, but SO42– reduction is no longer active within recent sapropels
(Passier et al., 1996a). The sapropels are usually underlain by gray sediments,
the so-called “protosapropel” zone (Maldonado and Stanley, 1976). Iron and
S are also enriched in this zone, whereas Corg is not. Pyrite has formed below
sapropels, causing the gray color of the “protosapropel” zone, as a result of
diffusion of HS–out of the HS–source (sapropel) (Passier et al., 1996a). Thus
the protosapropel did not form before sapropel deposition, but as a result of
downward sulfidation during and shortly after sapropel deposition, in this con-
text “synsapropel” would be a better term.
MATERIAL AND METHODS
Boxcore UM26 (lat 33°23.6′N, long 25°0.9′E, water depth 2160 m) con-
taining sapropel S1 was recovered 200 km south of Crete, during the 1994
Palaeoflux cruise of RV Urania. A subcore of UM26 was subsampled at a res-
olution of 0.5 to 1 cm aboard ship inside a N2filled glovebox and stored un-
der N2in air-tight containers at 4 °C.
Pyrite and Corg contents were measured according to methods described
by Passier et al. (1996a). For pyrite sulfur isotope analysis (δ34Spyr) H2S,
evolved during Cr(II) reduction, was trapped in a Cd-acetate solution as CdS;
AgNO3was subsequently added; and Ag2S was reprecipitated. Ag2S precipi-
tates were measured for their stable S isotope composition (34S/32S) relative to
the Vienna-Canyon Diablo troilite (V-CDT) by means of C-irmMS. Six sub-
samples (positions in Fig. 1) of about 2 g of wet sediment were rinsed with
demineralized water, dried inside a desiccator in vacuum, gently crushed, and
Geology; June 1997; v. 25; no. 6; p. 519–522; 4 figures; 1 table. 519
Pyrite contents, microtextures, and sulfur isotopes in relation to
formation of the youngest eastern Mediterranean sapropel
Hilde F.Passier
Department of Geochemistry, Utrecht University, P.O. Box 80021, 3508 TA Utrecht, The Netherlands
Jack J.Middelburg
Netherlands Institute of Ecology, Centre for Estuarine and Coastal Ecology, Yerseke, The Netherlands
Gert J. de Lange
Department of Geochemistry, Utrecht University, P.O. Box 80021, 3508 TA Utrecht, The Netherlands
Michael E. Böttcher
Institute of Chemistry and Biology of the Marine Environment, Carl von Ossietzky University, Oldenburg, Germany
ABSTRACT
Pyrite within and below sapropels in the eastern Mediterranean is a result of microbial
SO42– reduction within the sapropel, and the subsequent reaction of sulfide (HS–) with detrital
Fe and Fe2+ diffusing upward from underlying sediments. Below the youngest Mediterranean
sapropel, S1, pyrite (as much as 281 µmol pyritic S/g) is mostly present as euhedral crystals,
whereas within the sapropel only framboidal pyrite (as much as 360 µmol pyritic S/g) has been
detected. Framboidal microtextures indicate pyrite formation at the site of HS–production
within the sapropel. Euhedral pyrite, below the sapropel, forms when sulfate reduction in the
sapropel outcompetes iron liberation and supply, and HS–diffuses out of the sapropel. Sulfur
isotope values of pyrite are extremely light in the sapropel (–37.3‰ to –38.2‰) as well as below
the sapropel (–45.6‰ and –49.6‰), indicating that HS–has formed in a system with abundant
SO42– and in the presence of oxidants.
molded in resin. Thin sections of this resin were examined, and pyrite identi-
fied and photographed, with an electron microprobe–scanning electron mi-
croscope (JEOL 8600) and a CamScan series 4 scanning electron microscope.
PYRITE CONTENT AND IRON AVAILABILITY
The most important S species in sapropel S1 is pyrite as follows from the
comparison between Stot and Spyr (Fig. 1). Detailed analysis of other S species,
such as acid volatile sulfide, organic S, and elemental S, revealed that these
fractions were insignificant compared to pyrite (Passier et al., 1996b). Pyrite
and Corg are enriched in the sapropel. Pyrite is also enriched in a zone below
the sapropel (Fig. 1). The pyrite below the sapropel is formed from HS–that
diffused downward out of the sapropel during sapropel deposition (Fig. 2B;
Passier et al., 1996a). The pyrite maximum, however, is situated inside the
sapropel (Fig. 1); this means that Fe has been relatively enriched in the sapro-
pel. Thus, when the majority of the pyrite inside the sapropel formed
(Fig. 2A), HS–did not diffuse down out of the sapropel, but Fe2+ could diffuse
into the sapropel from below and hence result in enhanced pyrite formation
(Fig. 2A). This happens when SO42– reduction rates are low relative to the rate
of Fe liberation, as in the iron-dominated system described by Berner (1969).
The band-like enrichment of pyrite emphasizes the importance of pyrite for-
mation in the sediment.
The HS–source for pyrite formation is bacterial SO42– reduction inside
the sapropel. The Fe source comprises two pools: detrital Fe within the sapro-
pel and Fe diffusing from underlying sediments (Passier et al., 1996a, 1996b).
The amount of pyritic Fe within and below the sapropel minus a constant
amount of reactive detrital iron can be used to estimate the input of upward-
diffusing Fe during sapropel deposition (note that the pyrite present in the up-
per, now oxidized, part of the sapropel, and that present in the unrecovered
520 GEOLOGY, June 1997
Figure 1. Content vs. depth profiles of organic carbon (Corg), pyritic
sulfur (Spyr), and total sulfur (Stot). Pyritic sulfur normalized to the Al
content (Spyr/Al) shows that relative pyrite enrichment is not due to
any dilution effect.Triangles indicate positions of samples that were
studied with electron microscopy.The sapropel is indicated by the
shaded area.
Figure 2. Schematic representation of inferred pyrite formation condi-
tions and concentration gradients of HS–and Fe2+ (A) during periods
of relative low rates of sulfide production and framboidal pyrite forma-
tion within the sapropel and (B) during periods of relatively high rates
of sulfide production and euhedral pyrite formation below the sapropel.
part of the S-enriched zone below the sapropel are included). This flux is 0.07
µmol cm-2 yr-1, which can be compared to the upward Fe flux of 0.089
µmol cm-2 yr-1 in sapropel S1 at site UM26, as estimated from the pore-water
profile (van Santvoort et al., 1996). These two fluxes are in reasonably close
agreement, which suggests that the described scenario is plausible, and that
pyrite formation and/or Fe oxidation were able to act as traps for upward-dif-
fusing Fe2+.
PYRITE MICROTEXTURES
Within sapropel S1 pyrite occurs as framboids that have mean diameters
of 5 to 10 µm and consist of cubic microcrystals that are smaller than 1 µm
(Table 1, Fig. 3). Most pyrite within the sapropel is associated with shells; the
dimensions of the shells clearly influenced the shape of the framboids (Fig. 3).
Below the sapropel, however, pyrite is mostly present as euhedral (octahedral
and cubic) crystals with diameters of 2 to 10 µm and clusters (often with hex-
agonal crosscuts) of euhedral crystals (Table 1, Fig. 4). Furthermore, massive
pyrite shell infillings and large irregular pyrite aggregates are present below
the sapropel, and there is also pyrite which is not associated with shells. In the
intermediate sample DD01301, both framboids consisting of microcrystals
<1 µm and framboidal aggregates of euhedral crystals as much as 5 µm are
found (Table 1). Bioturbation probably occurred during the onset of sapropel
deposition, so that both organic-rich sapropel sediment and organic-poor sed-
iment are present in sample DD01301.
The size distribution of the framboids in sapropel S1 (5–10 µm and oc-
casionally larger) is typical for formation of pyrite in sediments near the sedi-
ment-water interface below an oxic or dysoxic water column (Wilkin et al.,
1996).
Our observations of framboidal pyrite in the sapropel and euhedral pyrite
below the sapropel are consistent with the mechanism proposed by Raiswell
(1982) and experimental results of Sweeney and Kaplan (1973),Rickard
(1997),Wang and Morse (1996). During the deposition of sapropel S1, Fe was
supplied by liberation from in situ sources and diffusion from suboxic sedi-
ments directly underlying the sapropel. As long as the supply of Fe could keep
pace with the production of HS–, HS–reacted in situ with Fe (i.e., at the site of
SO42– reduction), where HS–is continously produced. This situation resulted
in rapid formation of framboidal pyrite (Fig. 2A). At times of increased HS–
production relative to Fe liberation and/or supply, HS–diffused out of the Corg-
rich layer, titrating reactive Fe present below the sapropel and reacting with
upward-diffusing Fe2+ to form pyrite (Passier et al., 1996a). Below sapropel
S1, at the site of pyrite formation, the concentration of HS–was never as high
as in the sapropel (the HS–source) itself; the iron was less reactive, because
the most reactive Fe oxides had been reduced before; and slow diffusive pyrite
formation occurred. This led to the growth of euhedral pyrite below the sapro-
pel (Fig. 2B).
STABLE SULFUR ISOTOPES
The δ34S values of pyrite (δ34Spyr) in UM26 are –37.3‰ to –38.2‰ in
the sapropel, and –45.6‰ and –49.6‰ below the sapropel (Table 1). The
δ34Spyr reflects δ34S of sulfide formed, because limited fractionation occurs
when reduced S is transformed into pyrite (Price and Shieh, 1979). Mediter-
ranean seawater SO42– has a δ34S value of +20.6‰ (de Lange et al., 1990),
thus fractionations of 57.9‰ to 70.2‰ occurred. These fractionations are very
high: maximum isotopic fractionations of 40‰ to 70‰ between SO42– and
reduced S are usually observed in sediments (e.g., Goldhaber and Kaplan,
1974; Chambers and Trudinger, 1979). The observed fractionations of 57.9‰
to 70.2‰ imply that SO42– reduction took place in an open system, where ex-
change with seawater readily occurred, and thus no depletion of the SO42–
pool arose. Depletion would cause the remaining SO42– (and the reduced S
formed from it) to become progressively heavier (“reservoir effect,” e.g., Jør-
gensen, 1979). This exchange with seawater is possible close to the sediment-
water interface during sapropel deposition. However, the maximum fraction-
ation between sulfide and SO42– observed in laboratory studies with SO42–
respiring bacteria is 4‰ to 46‰, where the highest fractionations occurred at
the lowest SO42– reduction rates (e.g., Chambers and Trudinger, 1979). Fur-
ther fractionation has been reported to occur during bacterial HS–oxidation to
elemental S (or other intermediate oxidized S species such as thiosulfate), fol-
lowed by bacterial disproportionation (Canfield and Thamdrup, 1994). Ex-
tensive, repeated HS–oxidation results in lighter isotopic values (δ34S values)
of reduced S. If reoxidation invoked the high fractionations, oxidants must
have been able to reach the HS–that was formed in the sapropel. This may
have been nitrate and oxygen from the bottom water or sedimentary Fe and
Mn oxides. At times that HS–production is high relative to Fe liberation and
supply (Fig. 2B), HS–will diffuse downward and sulfidize sediments below
the sapropel. HS–may also move into the bottom waters, diffuse upward, and
reach the chemocline, where reoxidation occurs. Subsequently, evolved ele-
mental S settles back to the sediment and disproportionates, inducing more
negative S isotope values in the sediment. A similar scenario was proposed for
the present Black Sea (Canfield and Thamdrup, 1994). The HS–produced
during these periods of Fe limitation in the sapropel (Fig. 2B), will be incor-
porated into pyrite forming below the sapropel. When less HS–is incorporated
in pyrite at the site of HS–production, i.e., the sapropel, more HS–can partic-
ipate in the reoxidation cycle, and δ34S of pyrite forming below the sapropel
GEOLOGY, June 1997 521
Figure 3. Microprobe photo of framboidal pyrite (bright minerals) with
microcrystals <1 µm in sample DD01294 within sapropel. Pyrite is sit-
uated inside shell; growth is influenced by geometry of shell.
Figure 4. Scanning electron microscope photo of euhedral pyrite
(bright minerals) in sample DD01303, below sapropel; individual
cubes as well as clusters with hexagonal crosscuts are present.
will become even more negative. This might explain that δ34Spyr values below
the sapropel are more negative than within the sapropel. Moreover, pyrite
within the sapropel might also have been overprinted by relatively heavy re-
duced S that formed from pore-water SO42– after burial of the sapropel due to
a reservoir effect.
Note that part of the pyrite is formed in the absence of external oxidants
at the site of pyrite formation. Most pyrite reaction mechanisms include an ex-
ternal oxidant; however, direct pyrite formation with sulfide itself acting as an
oxidant has recently been made plausible (e.g., Drobner et al., 1990; Rick-
ard,1997), but the consequences for S isotope fractionation are not yet known.
IMPLICATIONS
The complementary information obtained from pyrite contents, micro-
textures, and sulfur isotopes defines the conditions of pyrite formation within
and below sapropel S1 in the eastern Mediterranean. All observed variations
in the properties of pyrite can be related to fluctuations in the production of
HS–relative to the supply of reactive Fe. When HS–production in the sapro-
pel is relatively low, Fe reaches the source of HS–inside the sapropel, a pyrite
enrichment forms, and pyrite formation is rapid, resulting in a framboidal tex-
ture. When HS– production is relatively high, HS–breaks out of the sapropel,
and forms the euhedral pyrite below the sapropel, in the so-called “synsapro-
pel.” During periods of high HS–production, bottom waters may eventually
become sulfidic. Nevertheless, on the basis of the size classification of fram-
boidal pyrite of Wilkin et al. (1996) the water column has been predominantly
oxic to dysoxic during sapropel formation.
As opposed to the system described by Raiswell (1982), the δ34S of the
euhedra is not heavier than the δ34S value of the framboids, because the source
of HS–is the open SO42– reduction system close to the sediment-water inter-
face for both textures, and there is no distinction between early diagenetic
pyrite formed close to the sediment-water interface and pyrite formed during
later diagenesis far below the sediment-water interface.
ACKNOWLEDGMENTS
Supported by the Cosiglio Nazionale della Ricerche, the Netherlands Orga-
nization of Scientific Research and the EU MAST-2 programme (CT93-0051). We
thank the crew and scientific parties of the Urania 1994 cruise, and H. de Waard,
D. van de Meent, G. Nobbe, T. Bouten, T. Broer, and P. Anten for analytical assis-
tance. C. H. van der Weijden is thanked for critically reading the manuscript. Re-
viewers R. Raiswell and R. Wilkin are thanked for their constructive remarks. Pub-
lication no. 970128 of the Netherlands School of Sedimentary Geology and no.
2245 of the Centre for Estuarine and Coastal Ecology,Yerseke, The Netherlands.
REFERENCES CITED
Berner, R. A., 1969, Migration of iron and sulfur within anaerobic sediments during
early diagenesis: American Journal of Science, v. 267, p. 19–42.
Calvert, S. E., 1983, Geochemistry of Pleistocene sapropels and associated sedi-
ments from the Eastern Mediterranean: Oceanologica Acta, v. 6, p. 255–267.
Canfield, D. E., and Thamdrup, B., 1994, The production of 34S-depleted sulfide
during bacterial dispropotionation of elemental sulfur: Science, v. 266,
p. 1973–1975.
Chambers, L. A., and Trudinger, P. A., 1979, Microbiological fractionation of stable
sulfur isotopes: A review and critique: Geomicrobiology Journal, v. 1,
p. 249–293.
de Lange, G. J., Middelburg, J. J., van der Weijden, C. H., Catalano, G., Luther,
G. W., III, Hydes, D. J., Woittiez, J. R. W., and Klinkhammer, G. P., 1990,
Composition of anoxic hypersaline brines in the Tyro and Bannock Basins,
eastern Mediterranean: Marine Chemistry, v. 31, p. 63–88.
Drobner, E., Huber, H., Wächtershäuser, G., Rose, D., and Stetter K.O., 1990, Pyrite
formation linked with hydrogen evolution under anaerobic conditions: Nature,
v. 346, p. 742–744.
Goldhaber, M. B., and Kaplan, I. R., 1974, The sulfur cycle, in Goldberg, E. D., ed.,
The Sea, Volume 5: New York, Wiley, p. 569–655.
Jørgensen, B. B., 1979, A theoretical model of the stable sulfur isotope distribution
in marine sediments: Geochimica et Cosmochimica Acta, v. 43, p. 363–374.
Maldonado, A., and Stanley, D. J., 1976, The Nile Cone: Submarine fan develop-
ment by cyclic sedimentation: Marine Geology, v. 20, p. 27–40.
Passier, H. F., Middelburg, J. J., de Lange, G. J., and van Os, B. J. H., 1996a, Dia-
genetic pyritisation under eastern Mediterranean sapropels caused by down-
ward sulphide diffusion: Geochimica et Cosmochimica Acta, v. 60,
p. 751–763.
Passier, H. F., de Lange, G. J., Middelburg, J. J., and van Os, B. J. H., 1996b, Sul-
phur appearances in and around sapropels, eastern Mediterranean, in Bottrell,
S. H., et al., eds., Proceedings of the Fourth International Symposium on the
Geochemistry of the Earth’s Surface: Leeds, United Kingdom, University of
Leeds, p. 101–104.
Price, F. T., and Shieh,Y. N., 1979, Fractionation of sulfur isotopes during labora-
tory synthesis of pyrite at low temperatures: Chemical Geology, v. 27,
p. 245–253.
Raiswell, R., 1982, Pyrite texture, isotopic composition and the availability of iron:
American Journal of Science, v. 82, p. 1244–1263.
Rickard, D., 1997, Kinetics of pyrite formation by the H2S oxidation of iron (II)
monosulfide in aqueous solutions between 25 and 125 °C: The rate equation:
Geochimica et Cosmochimica Acta, v. 61, p. (in press).
Rossignol-Strick, M., Nesterhoff, W., Olive, P., and Vergnaud-Grazzini, C., 1982,
After the deluge: Mediterranean stagnation and sapropel formation: Nature,
v. 295, p. 105–110.
Sweeney, R. E., and Kaplan, I. R., 1973, Pyrite framboid formation: Laboratory syn-
thesis and marine sediments: Economic Geology, v. 68, p. 618–634.
van Santvoort, P. J. M., de Lange, G. J., Thomson, J., Cussen, H., Wilson, T. R. S.,
Krom, M. D., and Ströhle, K., 1996, Active post-depositional oxidation of the
most recent sapropel (S1) in sediments of the eastern Mediterranean:
Geochimica et Cosmochimica Acta, v. 60, p. 4007–4024.
Wang, Q., and Morse, J. W., 1996, Pyrite formation under conditions approximating
those in anoxic sediments I. Pathway and morphology: Marine Chemistry,
v. 52, p. 99–121.
Wilkin, R. T., Barnes, H. L., and Brantley, S. L., 1996, The size distribution of fram-
boidal pyrite in modern sediments: An indicator of redox conditions:
Geochimica et Cosmochimica Acta, v. 60, p. 3897–3912.
Manuscript received December 2, 1996
Revised manuscript received February 21, 1997
Manuscript accepted March 12, 1997
522 Printed in U.S.A. GEOLOGY, June 1997
