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Phanerozoic marine phosphorites have been studied extensively
over the past decades, however the origin of these deposits is still a matter
of debate (e.g., Follmi, 1996; Trappe, 1998; Kouchinsky et al., 1999; Schulz
and Schulz, 2005; Dornbos, 2011; Bailey et al., 2013, Cosmidis et al., 2013,
Hiatt et al., 2015; Salama et al., 2015). Phosphorus and organic sedimentation in
the modern marine phosphorites have been supplied by upwelling currents along
continental margins of the western coast of Peru, Mexico and Namibia (e.g.,
Baturin, 1982; Jahnke et al., 1983; Glenn & Arthur, 1988; Follmi, 1996 and
references therein). However, phosphate-rich sediments can also be formed without
upwelling currents (e.g., Heggie et al., 1990; O’brien et al., 1990).
Hydrothermal processes and submarine volcanic activity have been considered as an
additional source of phosphorus for sedimentary phosphate enrichment (Berner,
1973; Froelich et al., 1982; Wheat et al., 1996; Petsch and Berner, 1998). Microbial
mediation processes in phosphogenic provinces may play an important role in phosphogenesis
and formation of phosphorites (e.g., Reimers et al., 1990; Krajewski et al.,
1994; Soudry, 2000; Schulz and Schulz, 2005; Bailey et al., 2007 and 2013; Cosmidis
et al., 2013; Lepland et al., 2014; Salama et al., 2015; Hiatt et al., 2015).

The formation of phosphorites involves mobilization and
redistribution of phosphate in surficial sediments under low sedimentation
rates, strong bottom currents, and an abundance of organic matter (e.g., Heggie
et al., 1990; O’brien et al., 1990; Follmi, 1996; Baturin, 1999). The decay of organic
matter releases phosphate to the pore waters, and contributes to supersaturation
that leads to phosphogenesis and precipitation of carbonate fluorapatite during
early diagenesis (Froelich et al., 1988; Glenn, 1990; Baturin, 1999). Marine phosphorites are divided into
two main types that record a wide spectrum of depositional and diagenetic
changes: pristine phosphorites which mostly represent the initial stages of
phosphogenesis and possess low P contents (e.g., Glenn & Arthur, 1988;
Piper et al., 1988; Föllmi, 1989; Soudry et al., 2013) and granular
phosphorites which reflect several episodes of superimposed phosphatization,
reworking by strong bottom currents and diagenesis in bottom sediments (Jarvis,
1992; Hiatt & Budd, 2003; Pufahl et al., 2003). Granular phosphorites are
more common in the ancient Phanerozoic record than pristine phosphorites, which
have usually been intensely reworked during deposition to form granular
phosphates (e.g., Soudry et al., 2013). Granular phosphorites in the ancient
sedimentary record consist of a variety of phosphatic particles including
phosphatized skeletal (i.e., fish bones and bioclasts) and non-skeletal
(nodules, coprolites, ovoids, peloids, coated grains, aggregates, and
lithoclasts) grains (e.g., Slansky, 1986). The enrichment of phosphates proceeds
in the granular phosphorites during the wanning stages between the repeated
cycles of reworking and redeposition of the phosphatic grains (e.g., Follmi,
1996; Baturin, 1999; Dornbos et al., 2006; Follmi et al., 2007; Dickinson and Wallace,

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The sedimentary Phanerozoic phosphorites of Iran were
developed in six periods: (1) lower Cambrian, (2) Ordovician-Silurian, (3)
upper Devonian, (4) upper Cretaceous, (5) Paleocene, and (6) Eocene-Oligocene
(Orris et al., 2015; Abedini and Asghar Calagari, 2017). These phosphorites are
distributed within five structural zones: (1) Central Iran, (2) Alborz-Azarbaidjan,
(3) Gorgan-Rasht, (4) folded Zagros, and (5) Sanandaj-Sirjan (e.g., Abedini and
Asghar Calagari, 2017). Devonian phosphorites are rare, in comparison
with the extensive and economic Phanerozoic phosphorites (Bentor, 1980). The most globally important Devonian phosphorite
deposits are found in Iran and Armenia (Halalat and Bolourchi,
1994). Other minor occurrences include thin sandy and pebbly phosphatic horizons
in the Plymouth Sound in UK (Humphreys and
Smith, 1989), Indiana and western New York, USA (Bluck, 1966; Baird, 1978).

The upper Devonian Geirud Formation is widely exposed
in the northern and central Iran and host to numerous sedimentary phosphatic
horizons. Extensive stratigraphic, sedimentological, and paleontological
investigations were carried out in the Geirud Formation in an attempt to
determine the depositional environments and paleoenvironmental conditions
dominated during the Devonian (e.g., Ghavidel?Syooki, 1995; Sharafi
et al. 2014, 2016). Comprehensive
studies concerning diagenetic evolution of the Geirud phosphorites have not
been done. Therefore, this study focuses mainly on integrating petrographic,
mineralogical and geochemical results to improve our understanding of the
post-depositional shallow to deep burial diagenetic evolution and the
subsequent subaerial weathering of the Geirud phosphorites, Northern Iran. The
work also aims at studying the distribution patterns of rare earth elements,
variations of Eu and Ce anomalies, and elemental ratios as useful means for
predicting the paleo-seawater conditions and diagenetic environments during
phosphatization of the upper Devonian Geirud phosphorites.


The Alborz Mountain Range (Figs. 1A, B) in northern
Iran extends for ~2000 km from the Caucasus Mountains and Azerbaijan in the
northwest to northern Afghanistan in the east. The stratigraphic succession of the
Alborz Mountain ranges from Precambrian to Holocene. The Palaeozoic sections in
Alborz start, from the base, with the Neoproterozoic-Lower Cambrian Soltanieh
Formation (shales and dolostones), the Cambrian Barut Formation (limestones,
siltstones, shales and micaceous sandstones with several thin volcanic
intercalations), and reddish brown siliciclastics of the Cambrian Zaigoon
(siltstones, shales), and Lalun (sandstones) formations.

The Lower Cambrian succession is overlain by the middle
Cambrian to early Ordovician Mila Forma (limestones members 1 to 4) and the early
to middle Ordovician brachiopod-bearing siltstones and shales member 5 of the
Mila and Lashkarak formations. The Silurian succession is represented by
shales, limestones, and sandstones of the Niur Formation that appears only in
the eastern part of the Alborz Mountain.

The Devonian sequence begins with intercalated
sandstones, shales, volcaniclastics and lava flows, overlain unconformably by
the Upper Devonian to Carboniferous siliciclastics and limestones of the
Khoshyelagh and Geirud formations and the Carboniferous limestone of the Mobarak
Formation. The Mobarak Formation is overlain by the Lower Permian sandstones, conglomerates,
and argillites of the Dorud Formation, limestones of the Middle Permian Ruteh
Formation and shales and limestones of the Upper Permian Nessen Formation.

The Alborz and the adjacent central part of Iran represent
remnants of the early Palaeozoic passive margin of Gondwana, which underwent an
important rifting phase during the Ordovician to Silurian (e.g., Stöcklin,
1968; Saidi & Akbarpour, 1992). This rifting phase was followed by extensive
continental shelf deposition from the middle Devonian to the middle Triassic, which
was associated with the Hercynian tectonic event (Ghorbani, 2013). During this
time span, Iran was located in the northern margin of Gondwana, along the southern
border of the Palaeo?Tethys Ocean at a latitude close to 30°S (e.g., Scotese
and McKerrow, 1990; Golonka, 2007; Bagheri and Stamfli, 2008). The
Alborz?Central Iran block collided with Eurasia, generating siliciclastic sediments
in continental to transitional zones and carbonates in shallow?marine to deeper
marine settings (Gaetani, 1965; Assereto, 1966; Bozorgnia, 1973; Stöcklin,
1974; Clark et al., 1975; Berberian & King, 1981; Alavi?Naini, 1993;
Alavi, 1996; Lasemi, 2001; Stampfli et al., 2001).

The Late Devonian (Frasnian?Famennian) age of the Geirud
Formation is well constrained on the basis of brachiopod (Bozorgnia, 1964),
palynomorph (Ghavidel?Syooki, 1995), trilobites (Wendt et al., 2005) and
goniatite (Dashtban, 1995) fossil remains.

of the Geirud Formation

Upper Devonian Geirud Formation in Alborz Mountain Range has numerous 0.5-7 m
thick phosphate horizons, which are intercalated with sandstones, black shales,
ferruginous carbonate strata and volcanic rocks (Assereto, 1963; Salehi, 1989; Halalat and
Bolourchi, 1994; Ghorbani, 2013).

The Geirud Formation crops
out at several localities in East Central Alborz such as QaemShahr, FiruzKuh, Gaduk,
Shahrud, PaQaleh and Damghan (Fig. 1C). The phosphate
resources at FiruzKuh and Gaduk are 56 Mt and 30.5 Mt
at 12 % P2O5, respectively (Orris
et al., 2015). At PaQaleh, the phosphate resources is 23 Mt at 9.42 % P2O5 (Orris
et al., 2015). In these areas, the Geirud Formation
disconformably overlies the marine shales of the Ordovician Milla Formation and
is conformably overlain by the dark-gray shales and limestones of the Mobarak
Formation (Halalat and Bolourchi 1994). The Geirud Formation is 240–370 m in thickness (from west to east) and was
deposited in a continental to shallow marine environments (Assereto 1966;
Bozorgnia 1973; Lasemi 2001; Alavi-Naini 1993 Weddige 1984; Gaetani 1965).

Detailed sedimentological and paleontological analyses
showed that Geirud Formation is organized into four main facies associations
that record sedimentation of a mixed?energy, wave? and tide?dominated estuary
passing seaward to an open?marine nearshore shelf environments (e.g., Ghavidel
Syooki, 1994; Sharafi et al., 2016). These facies associations include (i)
fluvial?dominated, bay?head deltaic deposits; (ii) estuarine deposits; (iii)
siliciclastic offshore?transition and shoreface sandstone deposits; and (iv)
carbonate open?marine shelf deposits (Sharafi et al., 2014, 2016). Moreover,
the presence of acritarchs in the Geirud Formation suggests a shallow marine environment
that was dominated in the Central Alborz Range during the upper Devonian
(Ghavidel Syooki, 1994). However, the presence of microspores suggests that
spore-bearing plants grew in the hinterland in a tropical climate and were
subsequently transported to a shallow marine basin (Ghavidel Syooki, 1994).

Geirud Formation at PaQaleh, north of Rudbar is about 80 m thick (Figs. 1C and
2A, B). It is composed of lower siliciclastic member of interbedded conglomerate,
phosphatic and non-phosphatic sandstones and shales along with subordinate phosphatic
horizons that may vary from 0.5 to 2 m in thickness (Fig. 2A, B). The upper
member of the Geirud Formation at PaQaleh consists of dolomitic limestones and
black shales with interbeds of black granular phosphorites (Fig. 2C). The
Geirud Formation is underlain by basalt and overlain conformably by the
Carboniferous black limestones of the Mobarak Formation (Fig. 2A, B).

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