New data from a continuously cored succession, the Mossy Grove core, near Jackson, central Mississippi, recovered ∼137 m of
marine clays (Yazoo Formation), spanning ∼5 Ma and including
the critical Eocene–Oligocene transition (EOT) event. These clay-rich
sediments yield well-preserved calcareous microfossil and palynomorph
assemblages. Here, we present a new organic-walled dinoflagellate cyst
(dinocyst) biostratigraphic framework, including the recognition of
23 dinocyst bioevents. These are integrated with new age
constraints based on calcareous nannofossil biostratigraphy and a
reassessment of the existing radiometric dates and planktonic foraminiferal
biostratigraphy, permitting the establishment of a robust and significantly
refined age model for the core. According to this new age model, a major
increase in sedimentation rate – from ∼2.1 to
∼4.7 cm kyr-1 – is observed at a core depth of
∼89.1 m (∼34.4 Ma). In the new age model the
section is significantly older than previously thought, by up to 1 Ma, with
the Eocene-Oligocene boundary (∼33.89 Ma) placed
∼34 m below the level previously identified. With these more
accurate age estimates, future isotopic and palaeoecological work on this
core can be more precisely integrated with other, globally distributed
records of the EOT.
Introduction
The Eocene–Oligocene transition (EOT: ∼34.0 Ma;
Westerhold et al., 2014)
represents a phase of accelerated climatic and biotic change that began
before and ended after the Eocene-Oligocene boundary (EOB: 33.89 Ma). It is
one of the most significant changes in the long-term background climate
state over the past 110 Ma (Friedrich et
al., 2012; Zachos et al., 2001). Across the whole EOT there is a
∼1.2–1.5 ‰ increase in benthic δ18O (e.g. Kennett
and Shackleton, 1976; Kennett, 1977; Miller and Mountain, 1987; Zachos et
al., 1996, 1999, 2001; Lear et al., 2000; Coxall et al., 2005), interpreted
to represent ice expansion equivalent to 60 % to 110 % of the present
Antarctic ice-sheet volume (Cramer et al., 2011) and a
3–4 ∘C deep-water cooling
(Cramer et al., 2011;
Lear et al., 2003). This ice expansion resulted in ∼50–60 m
of eustatic sea level fall
(Cramer et al.,
2011; Pekar et al., 2002), and it is associated with an estimated 3–6 ∘C cooling across latitudes
(Lear et al., 2008; Liu et al., 2009).
The presence of distinct rapid positive steps in high-resolution benthic
foraminiferal oxygen isotope (δ18O) records are interpreted as
representing at least two pulses of accelerated ice-sheet growth
(Coxall
et al., 2005; Scher et al., 2011). These steps appear to be orbitally paced,
with a duration of 40 kyr, separated by a 300–400 kyr interval of relative
stability, or plateau
(Coxall
et al., 2005; Scher et al., 2011; Westerhold et al., 2014). The first step
in δ18O records is thought to be associated with relatively
limited ice-sheet growth
(Katz et al.,
2008; Lear et al., 2008), being mainly driven by significant global cooling
(Houben
et al., 2018; Katz et al., 2008; Miller et al., 2008; Wade et al., 2012).
The second step in δ18O is thought to be dominated by
ice-volume expansion
(Katz et al.,
2008; Lear et al., 2008), a sharp deepening of the ocean carbonate
saturation depth (Coxall et al., 2005), and a
massive and abrupt increase in weathering flux from East Antarctica
(Scher et al., 2011). The
geological rapidity of the EOT is, most likely, driven by the non-linear
dynamics of Antarctic ice-sheet growth in response to a gradual decline in
atmospheric greenhouse gas (carbon dioxide, CO2) forcing
(DeConto
et al., 2008; DeConto and Pollard, 2003; Pagani et al., 2011; Pearson et
al., 2009). These same positive feedbacks also resulted in a strong
hysteresis, with the climate state becoming locked into a long-term
icehouse state with large continental-scale ice sheets on Antarctica since
the earliest Oligocene
(Cramer et
al., 2011; Lear et al., 2003; Zachos et al., 2001). Furthermore, the EOT is
associated with biotic turnover in plants and animals across a range of
latitudes and environments (Prothero, 1994). Although not
comparable to the “big five” mass extinction events (Raup and
Sepkoski, 1982), detailed micro-palaeontological records show that
substantial extinction and ecological reorganization is closely coupled to
the EOT
(Cotton
and Pearson, 2012; Coxall and Pearson, 2007; Dunkley Jones et al., 2008;
Houben et al., 2013; Moore et al., 2015; Pearson et al., 2008, 2009). These
include the extinction of the planktonic foraminiferal Family Hantkeninidae,
which marks the EOB (Nocchi et
al., 1988; Pearson et al., 2008; Premoli Silva and Jenkins, 1993),
significant extinction in shallow-water benthic foraminifera
(Cotton and Pearson,
2012; Pearson et al., 2008), radiolaria (Kamikuri and Wade, 2012),
and community overturning in the calcareous phytoplankton
(Bordiga
et al., 2015; Dunkley Jones et al., 2008; Persico and Villa, 2004; Villa et
al., 2008).
Dinocysts have been widely used in biostratigraphic and palaeoenvironmental
studies
(e.g.
Bujak and Williams, 1985; Duxbury and Vieira, 2018; Kothe, 1990; Powell,
1992; Pross et al., 2010; Stover et al., 1996; Vieira et al., 2018),
including several detailed studies of the EOT (e.g.
Brinkhuis
and Biffi, 1993; Brinkhuis, 1994; Jaramillo and Oboh-Ikuenobe, 1999; Van
Mourik et al., 2001; Oboh-Ikuenobe and Jaramillo, 2003; Houben et al., 2012,
2013, 2018). A major compilation of Late Cretaceous–Neogene calibrated
dinocyst bioevents, from both Northern Hemisphere and Southern Hemisphere, demonstrates
a strong climatic control on the timing of bioevents, which are commonly
diachronous (Williams et al., 2004) . Dinocyst biostratigraphic
studies thus need to consider comparable realms in similar climatic zones
when selecting appropriate dinocyst bioevents
(Williams and Bujak, 1985). There are existing
dinocyst biostratigraphic studies of the Eocene–Oligocene from both the
Northern Hemisphere
(Brinkhuis
and Biffi, 1993; Bujak and Mudge, 1994; Egger et al., 2016; Mudge and Bujak,
1994; Śliwińska et al., 2012; Thomsen et al., 2012; Wilpshaar et
al., 1996) and the Southern Hemisphere (Bijl et al., 2018; Wilson, 1988).
At the Massignano Eocene–Oligocene Global Stratotype Section and Point
(GSSP), there are two successive influxes of cool-water high-latitude
organic-walled dinoflagellate cyst (dinocyst) species
(Brinkhuis and Biffi, 1993), the first of which
correlates directly with the EOB and the second with the onset of a more
severe cold episode and inferred sea level lowstand. In parallel,
quantitative analysis of dinocyst distribution patterns from the
“Massicore”, central Italy (Van Mourik and Brinkhuis, 2005),
revealed biotic turnover potentially associated with latitudinal and climatic
zone changes through the Eocene–Oligocene transition. In the Massicore,
dinocyst assemblages are interpreted to represent substantial cooling on the
first step, with little subsequent change on the second step
(Houben et al., 2012).
Records from the Antarctic Margin show decreasing dinocyst species diversity
through the Eocene and at the EOB (Mohr, 1990), interpreted as the
progressive development of cold surface waters. More recent dinocyst studies
(Houben et al., 2013) document a sudden regime
shift in zooplankton-phytoplankton interactions in the Southern Ocean
associated with the earliest Oligocene glaciation of Antarctica, likely
triggered by cooling, ice-sheet expansion, and sea-ice formation.
Eocene to Oligocene strata are well represented in outcrop and the shallow
subsurface across the Gulf Coastal Plain of Mississippi and Alabama, as
well as in the offshore systems of the northern Gulf of Mexico
(Hosman, 1996). In these areas, dinocyst assemblages have an
instrumental importance for site-to-site stratigraphic correlation and
biochronology (Jaramillo and
Oboh-Ikuenobe, 1999). Existing palynological work from five sites across
southern Mississippi and Alabama allowed the correlation of stratigraphic
sequences through the upper Eocene and the lower Oligocene
(Jaramillo and Oboh-Ikuenobe, 1999).
They observed a maximum flooding surface, estimated to be ∼100 kyr older than the EOB and a lower Oligocene lowstand system tract, most
likely correlated to the global eustatic sea level lowstand coincident with
the Earliest Oligocene Glacial Maximum (EOGM;
Jaramillo and Oboh-Ikuenobe, 1999).
However, this study is not well integrated into biostratigraphic schemes
based on the calcareous plankton or detailed oxygen isotope stratigraphy,
both of which are important in generating robust correlations of dinocyst
biohorizons into global records of the EOT. More recently, coupled sea level
and palaeotemperature records from the neritic succession of the Saint
Stephens Quarry (SSQ), Alabama (Houben et al., 2018), are
interpreted to show cooling, minor ice-sheet expansion, and
temperature-driven turnover in dinocysts in the first step, while the second
step is characterized by a significant hiatus, potentially associated with a
major sea level fall, with no meaningful change in palynological
assemblages. Here we seek to extend this work by adding new biostratigraphic
data from a site, the Mossy Grove core, Jackson, Mississippi, which
recovered a continuous stratigraphic succession through the uppermost Eocene
and into the lower Oligocene. This core recovered a complete sequence
through the marine clays of the Yazoo Formation, central Mississippi
(Dockery III et al., 1991). This clay yields
palynomorph assemblages with high abundance and diversity, as well as an
exceptional preservation. The Mossy Grove core is also substantially
expanded compared to the St. Stephens Quarry section, with sedimentation
rates estimated to be more than 10 times higher through the EOT (St.
Stephens Quarry ∼0.4 cm kyr-1; Mossy Grove ∼4.7 cm kyr-1). This current study includes illustrated descriptions of the
assemblages from the Mossy Grove core and dinocyst stratigraphic range
data for the purpose of establishing a more resolved regional
biostratigraphic scheme for the Gulf of Mexico and the US Gulf Coast.
While considerable attention has focused on EOT sections in southwestern
Alabama
(Jaramillo
and Oboh-Ikuenobe, 1999; Katz et al., 2008; Mancini, 1979; Miller et al.,
1993, 2008; Quaijtaal and Brinkhuis, 2012; Tew and Mancini, 1995; Wade et
al., 2012) and southeastern Mississippi
(Jaramillo
and Oboh-Ikuenobe, 1999; Quaijtaal and Brinkhuis, 2012; Tew and Mancini,
1995), where the Yazoo Formation is lithologically heterogenous and divided
into multiple constituent members, in central and western Mississippi the
Yazoo Formation is a thick and relatively homogeneous succession of
calcareous clays. For micro-palaeontological and palaeoenvironmental studies,
these successions have the potential to yield some of the best marine
records of the late Eocene in all of North America, if not globally. The
relative lack of study of these successions appears to stem from early
problems in determining a robust chronostratigraphic framework and the
failure to constrain the EOB within the upper Yazoo Formation
(Obradovich et al., 1993). The purpose of this study is
to provide a detailed dinocyst biostratigraphic framework through an
expanded EOT succession from the US Gulf Coast, as the basis for future
studies of community change, extinction, and palaeoenvironments based on these
assemblages. Such studies of the response of organic-walled cyst-producing
dinoflagellate communities across this interval are important for an
understanding of the nature of the reorganization of planktonic ecosystems
through this major transition.
Regional geological context
The Gulf Coastal Plain (Fig. 1) is one of the
largest physiographic provinces in North America. In the subsurface the
Mississippi embayment consists of a structurally complex basin with thick
Jurassic-to-Holocene deposits (Cushing et al., 1964;
Hosman, 1996). Within these sequences, the most continuous Paleogene
successions are found between central and southern Mississippi and Alabama
(Cushing et al., 1964; Hosman, 1996).
Palaeogeographic reconstruction showing location of the Mossy Grove, #1 Wayne and #1 Young cores, as well as St. Stephens Quarry outcrop. Main North American palaeogeography at ∼35 Ma follows Deep Time Maps™ (https://deeptimemaps.com/, last access: 16 October 2018).
The upper Eocene to lower Oligocene stratigraphy of the US Gulf Coast is
split into two groups (Fig. 2) – the Jackson
(∼ upper Eocene) and Vicksburg (∼ lower
Oligocene) groups (Cushing et al., 1964; Hosman, 1996).
The contact between these groups is commonly indistinguishable on a
lithologic basis, but they have distinct faunal assemblages
(Hosman, 1996). The lithology, and constituent formations within
these groups, also varies geographically across Mississippi and Alabama,
reflecting relative positions along the palaeoshelf and proximity to river
outflows. The upper middle Eocene of Mississippi consists of discontinuous
and lenticular beds of lignitic-to-carbonaceous fine-to-medium quartz sand,
silt, and clay. These sediments constitute the non-marine Cockfield
Formation, the youngest continental deposits of the Eocene on the Gulf
Coastal Plain (Cushing et al., 1964; Hosman, 1996). The
upper Eocene Jackson Group overlies the Cockfield Formation, and it represents
the last extensive marine inundation within the Mississippi embayment
(Cushing et al., 1964; Hosman, 1996). The Jackson Group
consists of two major facies: to the west of the Gulf Coastal Plain, towards
Texas, a shallow-water marine facies with beach sands, clays, and occasional
volcanic tuffs, whilst to the east, in the central part of the Gulf Coastal
Plain (Mississippi and Alabama), the group becomes more fossiliferous and
argillaceous as deeper-water clays dominate (Cushing et
al., 1964; Hosman, 1996). The two main lithostratigraphic units within the
Jackson Group are the highly fossiliferous glauconitic sandy marls of the
basal Moodys Branch Formation and the calcareous fossiliferous dark-grey to
blue clays of the Yazoo Formation (Cushing et al., 1964;
Hosman, 1996). The Yazoo Formation has four well-characterized members in
eastern Mississippi and Alabama: the North Twistwood Creek Member (marl and
calcareous clay), the Coccoa Sand Member (sandy marl), Pachuta Marl Member
(glauconitic fossiliferous marls and sandy hard limestones), and the Shubuta
Member (calcareous clays with concretions) (Cushing et
al., 1964; Hosman, 1996). In central Mississippi, in the region of the Mossy
Grove core, the entire Yazoo Formation is represented by an undifferentiated
marine clay unit – the Yazoo Formation
(Dockery III et al., 1991). Towards northern
Mississippi this group occurs as a non-marine clastic facies, with evidence
of lacustrine, palustrine, and lagoonal sedimentation, and the Yazoo
Formation becomes undifferentiated and relatively homogeneous
(Cushing et al., 1964; Hosman, 1996). The Oligocene
series begins with the marine fossiliferous glauconitic clay and sandy clay
of the Red Bluff Formation, overlain by the fossiliferous and laminated
sands of the Forest Hill Formation. Westward, this unit presents some lenses
of lignite and lignitic clay. Above the Forest Hill Formation is the
fossiliferous, granular, white crumbly limestone of the Bumpnose Formation
(Cushing et al., 1964; Hosman, 1996). The Bumpnose
Formation in Alabama and Florida is the equivalent of the Red Bluff–Forest
Hill sequence in central and southern Mississippi (Hosman,
1996). Finally, overlying all these units is the Marianna–Mint Spring
Formation, which consists of fossiliferous and porous limestone, glauconitic
marl, and calcareous clay (Cushing et al., 1964; Hosman,
1996). In northern Mississippi, no evidence of Oligocene sedimentation is
observed (Cushing et al., 1964; Hosman, 1996).
Simplified stratigraphic correlation chart of upper Eocene to lower Oligocene sedimentary rocks in the Mississippi–Alabama area (modified from Pasley and Hazel, 1995; Tew and Mancini, 1995; Hosman, 1996).
The upper Eocene to lower Oligocene lithostratigraphy outlined above has
been interpreted within two distinct sequence stratigraphic models. The
first argues that the Yazoo–Bumpnose contact coincides with a maximum
flooding surface throughout the region
(Baum and Vail,
1988; Echols et al., 2003; Jaramillo and Oboh-Ikuenobe, 1999; Loutit et al.,
1988; Mancini and Tew, 1991; Tew, 1992). The second, and more recent
(Miller et al.,
2008), supports the association of the Yazoo–Bumpnose contact
(Dockery, 1982) with a low stand sequence boundary, which is
linked to the increasing δ18O
(Zachos
et al., 2001, 1996, 2008) and global eustatic sea level fall
(Coxall
et al., 2005; Pekar et al., 2002) during the EOT. Sequence stratigraphic
interpretations are difficult because of the depositional and lithologic
variability across this region, and the difficulties in establishing robust,
mostly biostratigraphic, tie points between key sections
(Miller et al.,
2008). Some of the key sections, from both outcrops and cores, which have
been used for regional stratigraphic correlation, chronostratigraphy, and
sequence stratigraphy, are shown in Fig. 3.
Stratigraphic correlation between the Mossy Grove core and the other three sections from Mississippi and Alabama (Jaramillo and Oboh-Ikuenobe, 1999). Lithology at the Mossy Grove follows Dockery III et al. (1991) and the timescale for all sections follows Westerhold et al. (2014). See discussion about dinocyst bioevents in Sect. 3.2.
Even at the St. Stephens Quarry section, southwestern Alabama, which is a
widely studied reference for the EOT
(Katz
et al., 2008; Miller et al., 2008; Wade et al., 2012), establishing a robust
biostratigraphy and sequence stratigraphy has been difficult
(Miller et al.,
2008). The low abundance and generally poor preservation of hantkeninids in
the shallow-water Gulf Coast successions make the accurate placement of
their last occurrence (LO) problematic
(Miller et al.,
2008). This difficulty is further aggravated by the reworking of upper
Eocene calcareous microfossils up into lowest Oligocene sediments as a
result of the major regression across the EOT (Bybell and Poore,
1983). Nevertheless, a middle-to-late Priabonian age was assigned to the
Pachuta Member at SSQ, based on the LO of R. reticulata, upper part of nannofossil zone NP19-20 (Miller et al., 2008). The Shubuta Member has a late Priabonian age
correlated to the lower part of NP21 zone, the interval between E16 and the
lower part of O1 zones, and the upper part of magnetochron C13r
(Miller et al.,
2008). The Red Bluff–Bumpnose sequence was dated as late Priabonian to early
Rupelian age and can be correlated to the middle part of NP21 zone, the
lower to middle part of O1 zone and the upper part of C13n magnetochron
(Miller et al.,
2008).
Materials and methodsMossy Grove borehole
The Mossy Grove borehole was drilled in September 1991 at Mossy Grove, Hinds
County, Mississippi. It consists of a 161.6 m (530 ft) continuously cored
succession, with near full recovery (Dockery
III et al., 1991). At its base it recovered ∼5.5 m of the upper
Cockfield Formation (161.6 to 156.1 m; 530 to 512 ft, 18 ft thick), overlain
by ∼3.7 m of the Moodys Branch Formation (156.1 to 152.4 m;
512 to 500 ft, 12 ft thick), and then a thick 140.9 m succession of the
Yazoo Formation (152.4 to 11.6 m; 500 to 38 ft, 462 ft thick). At the top of
the core, overlaying the Yazoo Formation, 3.0 m of terrestrial lignites of
the Forest Hill Formation were recovered (11.6 to 8.5 m; 38 to 28 ft, 10 ft
thick), which in turn was overlain by a 8.5 m thick (28 ft) cover of
Pleistocene loess. To improve the chronology of the upper Yazoo Formation,
40Ar/39Ar absolute dating was undertaken from a number of
bentonite layers at four sites (Obradovich et al.,
1993; Priddy, 1960), including seven layers within the Mossy Grove core,
from which two dates were recovered (Obradovich et
al., 1993; Obradovich and Dockery III, 1996).
Establishing a robust age model for the Mossy Grove core has required
significant effort and the integration of dinocyst, calcareous nannofossil,
and radiometric dating techniques. Initial age models were based on
planktonic foraminifera assemblage data
(Fluegeman, 1996; Fluegeman et al.,
2009), but due to low planktonic foraminifera abundances, especially in the
upper part of the core (< 91.4 m; 300 ft), this age model relied
upon poorly calibrated secondary markers. For instance, the direct placement
of the EOB, based on the last occurrence of the Hantkeninidae, is
problematic as the record of Hantknenina alabamensis, the only hantkeninid species found in the
core, is poor and discontinuous above ∼86.6 m (284 ft). The
last occurrence of Turborotalia cerroazulensis, expected to be ∼0.1 Myr older than the
EOB and the Hantkeninidae extinction, occurs ∼7.3 m (24 ft)
higher than the last occurrence of H. alabamensis.
Radiometric 40Ar/39Ar dating is significantly offset from the
existing planktonic foraminifera biochronology. Bentonite layers at
∼26.5 m (87 ft) and ∼84.4 m (277 ft) were
dated to 33.40 and 34.36 Ma, respectively (Obradovich and Dockery
III, 1996), whereas foraminifera biochronology places these bentonites at
∼33.87 and 35.33 Ma, respectively
(Fluegeman et al., 2009), an offset of
between ∼0.5 and 1 Myr. This mismatch has implications for the
accurate placement of the EOB in this section; assuming a constant
sedimentation rate within the interval between the bentonites, radiometric
dating would place the boundary at ∼56.7 m (186 ft), which is
∼30.2 m (99 ft) lower than the existing foraminiferal
estimate of ∼26.5 m (87 ft)
(Fluegeman et al., 2009).
Palynological preparation techniques
Altogether, 112 sediment samples were collected from the Mossy Grove core at
∼1.2 m (4 ft) intervals, between ∼17 and 152 m
deep (55 and 499 ft). A mass of 20 g from each sediment sample was broken into pieces,
∼5 mm in diameter, using a pestle and mortar. This was
followed by a set of acid treatments, including (1) 40 % HCl for 30 min to remove carbonates; (2) 60 % HF for 24 h to disaggregate the
rock matrix; (3) sieving of the residues at 10 µm through a nylon mesh
sieve to retain the HF effluent from the material; (4) a second HCl treatment
to remove any precipitate formed during stage 2; and (5) a final sieving, as per
stage 3. The material was then subjected to oxidation (70 % HNO3 for
exactly 2 min) to remove pyrite, inorganic debris, and any unstructured
organic material from the residues. Another sieving step was done to remove
any HNO3 effluent. A final cleaning treatment was undertaken with a
combination of domestic and industrial detergents. Using a glass swirling
dish and centrifugal swirling techniques, palynomorphs in each sample were
subsequently concentrated and Bismark brown was added to make them more
visible with light microscopy. Finally, the samples were sieved into two
size fractions, 10–30 µm (concentrating spores and pollen) and
> 30 µm (concentrating dinocysts), and then mounted on
separate 22mm×22 mm coverslips, which were glued to a glass slide using
Norland optical adhesive. In this work, only the coarse fraction content of
each slide was analysed. A pilot survey of these slides revealed that the
acid and oxidizing technique yielded higher diversity than their non-acid
and non-oxidizing counterparts (Burgess, 2015).
All slides were stored in the collection of the School of Geography,
Earth and Environmental Sciences, University of Birmingham.
Dinocyst assemblage characterization
With a Zeiss transmitted-light microscope, at least 200 dinocyst specimens
were counted in each sample, along with the respective number of spores,
pollen, algae (Prasinophyceae and Chlorophyceae), zoomorphs/zooclasts,
phytoclasts, and amorphous organic matter. Once 200 specimens were counted,
the remaining fields of the coverslip were scanned for very rare dinocysts.
Only palynomorphs that were more than 50 % complete and not obscured
either by air bubbles or organic debris were considered. We used the
following abundance scheme: 0 % – barren (B); 0 %–1 % – trace (T);
1 %–10 % – few (F); 10 %–25 % – common (C); > 25 % –
abundant (A). Preservation was qualitatively categorized as good (G),
moderate (M) or poor (P).
Nannofossil assemblage characterization
Samples for the analysis of calcareous nannofossils were taken at
∼1.2 m (∼4 ft) intervals throughout the Mossy
Grove core. Samples were prepared using the simple smear slide technique
(Bown and Young, 1998) and observed with standard
transmitted cross-polarized light (XPL) microscopy (Zeiss AxioScope at 1250× magnification). Calcareous nannofossil preservation and abundances were
logged for all samples with taxonomy following
Dunkley Jones et al. (2009). Major
biohorizons were identified based on the presence and absence of key marker
taxa following the biostratigraphic scheme of
Agnini et al. (2014).
ResultsDinocyst distribution
Nearly all the core samples contain abundant and very well-preserved
dinocysts, with only a few samples being scarce or rare. A
semiquantitative distribution of all dinocysts recognized in the
investigated section is shown in Fig. 4.
Spiniferites species dominate in most samples and account for ∼50 % of
the overall count, followed by Hystrichokolpoma (mainly H. rigaudiae and H. salacia) and Operculodinium (mainly O. centrocarpum)
(∼6 % each); Charlesdowniea (C. coleothrypta) (∼4 %); Heteraulacacysta (mainly H. porosa);
Lingulodinium (mainly L. Machaerophorum) and Cleistosphaeridium (mainly C. ancyreum) (3 % each); and Cordosphaeridium, Dinopterygium (D. cladoides), Glaphyrocysta, Homotryblium (mainly H. floripes),
Saturnodinium (S. pansum), and Thalassiphora (mainly T. fenestrata) (∼2 % each). These genera account for
nearly 90 % of all dinocysts counted in the section, although there are
other significant occurrences of Achomosphaera alcicornu, Dapsilidinium pastielsii, Enneadocysta arcuata, and Hystrichosphaeridium tubiferum. Several influxes of relative
abundances higher than 25 % are observed: C. fibrospinosum (∼142.3 m depth;
467 ft), D. cladoides (∼94.8 m depth; 311 ft), H. leptalea (∼66.8 m
depth; 219 ft), H. porosa (∼58.4 m depth; 191.5 ft), H. floripes (∼118.0 m depth; 387 ft), H. tubiferum (∼91.1 m depth; 299 ft), L. machaerophorum
(∼107.0 m depth; 351ft), O. centrocarpum (∼68.0 and
∼75.3 m depth; 223 and 247 ft), T. fenestrata (∼83.8
and ∼85.0 m depth; 275 and 279 ft), and W. articulata (∼102.1 m depth; 335 ft).
Semiquantitative range chart of dinocyst taxa encountered at the Mossy Grove core, central Mississippi, US Gulf Coastal Plain, including calcareous nannoplankton zonation.
Nannofossil biostratigraphy
For this study we present the first calcareous nannofossil biostratigraphy
of the Mossy Grove core, with biohorizons calibrated to the global zonation
scheme of Agnini et al. (2014). We updated these absolute ages
(Agnini et al., 2014) to the
timescale of Westerhold et al. (2014). Three of these bioevents constrain the base and top of the core: the
absence of Sphenolithus obtusus at the base of the core indicates that the section bottom (152.1 m; 499 ft) is younger than 38.24 Ma; the first common occurrence (FCO) of Isthmolithus recurvus was identified at
∼136.9 m (449 ft), which corresponds to 36.74 Ma; and the LO
of Coccolithus formosus, whose occurrence spans from ∼50.0 m (164 ft) to
∼16.8 m (55 ft), indicates that the top of the section is
older than 32.90 Ma. Moreover, two other LO events have been detected within
the core: Discoaster saipanensis (∼89.2 m; 292.5 ft) and Reticulofenestra reticulata (∼107.6 m;
353 ft), indicating ages of 34.44 and 35.31 Ma, respectively. All
biostratigraphic and radiometric ages are presented in the
Table 1.
List of bioevents, biohorizons, and calibrated ages, updated to the timescale of Westerhold et al. (2014). Legend: CN – calcareous nannofossils (this study); PF – planktonic foraminifera (Fluegeman et al., 2009); RD – radiometric dating (Obradovich et al., 1993; Obradovich and Dockery III, 1996).
TypeEventDepth Age (Ma)mftCNLO C. formosus (constrain)16.855.032.90RD40Ar/39Ar26.587.033.40PFE16/O1 boundary29.095.033.82PFE15/E16 boundary56.4185.034.50RD40Ar/39Ar84.4277.034.36CNLO D. saipanensis89.2292.534.44PFE14/E15 boundary100.3329.035.90CNLO R. reticulata107.6353.035.31PFE13/E14 boundary124.1407.038.02CNFCO I. recurvus136.9449.036.74CNLO Sphenoliths obtusus (constrain)152.1499.038.24Dinocyst biostratigraphic framework
Dinocysts dominate the palynological assemblages in the studied material.
In total, we identified 52 genera and 70 species, with nomenclature
following Williams et al. (2017) and references therein. A
full species list is provided in Appendix 1. The rich and well-preserved
dinocyst assemblages provide the basis for a detailed assessment of upper
Eocene dinoflagellate biostratigraphy in this succession, which is developed
with reference to dinocyst biostratigraphic data from offshore eastern
Canada (Egger et al., 2016;
Williams, 1975, 1977), the Gulf Coast Plain
(Houben et al., 2018;
Jaramillo and Oboh-Ikuenobe, 1999), offshore Florida
(VanMourik et al., 2001), the Norwegian and Greenland
seas
(Eldrett
et al., 2004; Manum, 1976; Manum et al., 1989; Williams and Manum, 1999),
the North Sea
(Bujak
and Mudge, 1994; Gradstein et al., 1992; Hansen, 1977; Heilmann-Clausen and
Van Simaeys, 2005; Mudge and Bujak, 1996; Śliwińska et al., 2012),
the Hampshire Basin (Costa et al., 1976), the central
Mediterranean region
(Brinkhuis,
1994; Brinkhuis and Biffi, 1993; Van Mourik and Brinkhuis, 2005; Wilpshaar et
al., 1996), northwestern Europe (England, Belgium, and Germany;
Costa and Downie, 1979; Kothe, 1990), the Tasmanian Gateway
(Stickley et al., 2004), and the wider
Northern Hemisphere (Williams, 1993). Although there have been
studies of dinocyst biohorizons across the EOB interval in Gulf Coast
sections from Alabama (e.g.
Jaramillo and
Oboh-Ikuenobe, 1999; Houben et al., 2018), these do not have the
stratigraphic coverage of the upper Eocene or the quality of dinocyst
preservation provided by the Mossy Grove core material. Age assignments are
based on the global compilation of age-calibrated bioevents
(Williams et al., 2004), with a total of 23 potentially
useful dinocyst biohorizons recognized in the Mossy Grove succession
(Table 2), including 7 first occurrences (FO),
1 first common occurrence (FCO), 14 last occurrences (LO), and 1
last common occurrence (LCO) event.
List of dinocyst bioevents recorded in the section. Reliability rating and absolute ages were assigned for selected events. Timescale follows Westerhold et al. (2014). The last column presents ages calculated from the age–depth model provided by this study (see Sect. 3.4). Legend: FO represents first occurrence; LO represents last occurrence; FCO represents first common occurrence; LCO represents last common occurrence; A represents acme; * represents age available only after the timescale of Pälike et al. (2006); H represents high reliability; L represents low reliability; and M represents moderate reliability.
References for chosen ages: a – Brinkhuis and Biffi (1993), b – Brinkhuis et al. (2003), c – Bujak and Mudge (1994), d – Eldrett et al. (2004), e – Jaramillo and Oboh-Ikuenobe (1999), f – Williams et al. (2004), and g – Williams and Manum (1999).
At least 15 of the 23 biohorizons within the Mossy Grove
succession could be correlated to other mid-latitude areas in the Northern
Hemisphere, and we discuss each of these – from the base of the succession
upwards – in the text below. Our estimation of the reliability of each
bioevent as highly, moderately, or lowly reliable; it is based on a combination
of the clarity of taxonomic definition of a species or genus, as well as the
number of reports of the associated biohorizon at other localities, and the
consistency of stratigraphic position between them. The estimated age of
each bioevent is also given, along with a rationale for the chosen
calibration. Ages are first provided according to the original assignment
and timescale but then, for consistency, are all updated to the
Westerhold et al. (2014)
timescale, which is given in brackets. Where there is evidence for
latitudinal diachroneity of an event, we preferentially choose age
calibrations from locations closest to the study site. Standard
nannoplankton (Agnini et al.,
2014; Martini, 1971) and foraminiferal
(Berggren
et al., 1995; Berggren and Pearson, 2005; Wade et al., 2011) biozonation
schemes were used when discussing selected dinocyst bioevents.
The palynological assemblages of the Mossy Grove core show strong evidence
for sea level fall associated with the EOT interval manifest in the form of
increased flux of terrestrial plant material but also in the influx of
clearly weathered and reworked acritarchs and continental palynomorphs
towards the top of the succession. This includes reworked Late Cretaceous
taxa such as Achomosphaera ramulifera, Cannosphaeropsis sp., Chatangiella sp., Cymososphaeridium phoenix. Dinogymnium sibiricum, Litosphaeridium siphoniphorum, and Palaeohystrichophora infusorioides, as well as early and middle Eocene
dinocyst taxa Hystrichosphaeridium tubiferum (Bujak and Mudge,
1994; Mudge and Bujak, 1996), Wetzeliella articulata
(Bujak and Mudge, 1994; Mudge
and Bujak, 1996), Piladinium columna Bujak and
Mudge, 1994; Mudge and Bujak, 1996), and Diphyes ficusoides (Gradstein
et al., 1992). The distribution patterns of Eocene dinocysts through the
core indicate local reworking of older middle and lower Eocene marine
sediments from around the margin of the Mississippi embayment, especially
through the EOT and towards the top of the succession. This reworking makes
the confident placement of some last occurrences difficult, especially where
there is some existing discrepancy between published age assessments of
these bioevents as being either late Eocene or early Oligocene. Here we
interpret consistent presence of a taxa, followed by a continuous
stratigraphic absence or sporadic occurrences, as representing a genuine LO
event. Sporadic occurrences are reported as local reworking from exposed
middle to upper Eocene sediments during regional sea level fall.
FO Hemiplacophora semilunifera. The first occurrence of this species is at ∼143.6 m (471 ft). From this point upwards, it is present in most samples, with abundances
gradually increasing, especially in the interval ∼119.2–99.7 m (391–327 ft). Based on data from equatorial and Southern Hemisphere sites
(Stickley et al., 2004; Williams et
al., 2004), this event has been tentatively dated between 36.00 and 41.40 Ma
(36.05 to 41.20 Ma), in middle and higher latitude sites, respectively.
FO Lentinia serrata. This species first occurs at ∼133.8 m (439 ft), although
there is a short interval after this occurrence in which the taxon is not
observed. This event was identified in central Italy
(Brinkhuis and Biffi, 1993) sections in the
lowermost Adi interval zone, which corresponds to the calcareous
nannoplankton zone NP21, the planktonic foraminifera zone P18, and the top
of the magnetostratigraphic chron C13r. This agrees with an age of 33.64 Ma
(33.84 Ma) assigned in equatorial latitudes (Williams et al.,
2004). Although these data seem consistent, dating from the mid-latitudes in
the Northern Hemisphere (Williams et al., 2004) points to an
age of 40.00 Ma (39.97 Ma) for this bioevent.
LO Areoligera sentosa. The last occurrence of this species is clearly observed at
∼124.1 m (407 ft). This event is considered synchronous and
an excellent chronostratigraphic marker in the entire North Sea
Basin to Norwegian–Greenland seas region
(Bujak and Mudge,
1994; Costa et al., 1976; Heilmann-Clausen and Van Simaeys, 2005; Mudge and
Bujak, 1996), with an age assignment in the lower part of E7 dinocyst zone
and, more precisely, at the E7a/E7b subzone boundary
(Bujak and Mudge, 1994; Mudge
and Bujak, 1996), which identified this event just between NP16 and NP17
nannoplankton zones and in the lower part of planktonic foraminifera zone
P14.
LO Rottnestia borussica. Almost all occurrences of this species at the study site are below
∼137.5 m (451 ft), except for an isolated occurrence at
∼110.6 m (363 ft). Although this isolated appearance may
reflect patterns of ecological exclusion, interpreting it as LO event would
be less consistent with the age model. Therefore, we conclude that the
isolated upper occurrence represents reworking processes and constrain the
range of this taxa to a short interval in the lower part of the succession
(below 137.5 m). In the Hampshire Basin (Aubry, 1983, 1985) and
the North Sea (Bujak and Mudge,
1994; Mudge and Bujak, 1996) this event was dated as late Bartonian,
calibrated with the upper part of the dinocyst subzone E7b
(Bujak and Mudge, 1994; Mudge
and Bujak, 1996), nannoplankton zone NP17, and planktonic foraminifera zone
P14, whilst in Italy its range is reported to extend into the lower Rupelian
(Brinkhuis and Biffi, 1993). A late Bartonian age
of 39.00 Ma (39.00 Ma) is consistent within the Norwegian–Greenland seas
(Eldrett et al.,
2004; Eldrett and Harding, 2009) and with the southwestern margin of the
Rockall Plateau, North Atlantic, where it was dated as middle Eocene and
calibrated within the dinocyst zone within zone IVb (Costa
and Downie, 1979); however, it is distinct from middle Priabonian age assigned in
the central Danish basin (Heilmann-Clausen and Van Simaeys,
2005), calibrated with the middle part of the nannoplankton zone NP19/20; in offshore Florida (Van Mourik et al., 2001), it is
found within planktonic foraminifera zone P16. Based on correlation to
nannofossil biostratigraphy within the Mossy Grove core, we can constrain
this event to be younger than 39.00 Ma, such that this taxa ranges into the
topmost Bartonian.
LO Hemiplacophora semilunifera. The last occurrence of H. semilunifera is clearly identified at ∼100.9 m (331 ft), above which the taxa is absent, except for very rare isolated
and widely spaced occurrences. This event was dated in the Mediterranean
(Brinkhuis and Biffi, 1993; Wilpshaar
et al., 1996) with an early Oligocene age, correlated to the top of the Gse
dinocyst zone (Brinkhuis and Biffi, 1993), the
nannoplankton NP21, and foraminifera P18 biozones. Nevertheless, estimates
from middle latitudes in the Southern Hemisphere (Williams et
al., 2004) indicate that it occurred much earlier, during the NP19-20
biozone and between the P15 and P16 biozones, which an assigned age of 35.20 Ma (35.49 Ma).
FO Achomosphaera alcicornu. The first occurrence of this taxon at the study site was detected at
∼66.8 m (219 ft), and it is consistently observed until the
top of the section. Early estimates from Danish sections dated this bioevent
as Danian, placing it in the upper part of the S. inornata subzone (Hansen,
1977). In high latitudes of the Southern Hemisphere, this event was dated at
35.70 Ma (Williams et al., 2004; 35.82 Ma). However,
estimates from lower latitudes, such as the Mediterranean
(Brinkhuis and Biffi, 1993; Wilpshaar
et al., 1996) and the Equator (Williams et al., 2004),
identified this event within the nannoplankton NP21 and foraminifera P17
biozones, and within the magnetostratigraphic chron C13r, with an assigned
age varying between 33.76 (Williams et al., 2004; 33.99 Ma)
and 34.00 Ma (Wilpshaar et al., 1996; 34.29 Ma).
LO Diphyes colligerum. This species has a relatively continuous record throughout the entire
core up to ∼50.9 m (167 ft). After an interval of absence
(∼22.0 m; 72 ft), it reappears near the top of the hole
(∼29.0 m; 95 ft) associated with another species of the same
genus, D. ficusoides, which is known to have been extinct since the middle Lutetian. This
suggests that the D. colligerum specimens found near the top of the core are most likely
the result of reworking. Although the last occurrence of D. colligerum has been used to
indicate the Eocene–Oligocene boundary worldwide (Williams,
1993), other studies record this event from the mid-Lutetian to the early
Oligocene. In the southwestern margin of the Rockall Plateau, North
Atlantic, this event was dated as middle Eocene and identified within
zone IVb (Costa and Downie, 1979). In the North Sea it has been
identified close to the Lutetian–Bartonian boundary
(Bujak and Mudge, 1994; Mudge
and Bujak, 1996), similar to its position in the Norwegian–Greenland seas
(Eldrett et al., 2004), although
later studies found it ranging up into the middle Bartonian in this area
(Eldrett and Harding, 2009). In the central Danish
basin it corresponds to the middle Lutetian, in the upper part of NP15
nannoplankton biozone (Heilmann-Clausen and Van Simaeys,
2005). In northwest Germany it was recorded in several samples from the
lower Oligocene (Kothe, 1990). Across North America, however, some
consistent younger ages were assigned for this event, from late Priabonian
to early Rupelian, including offshore Canada (nannoplankton Biozone NP21)
(Williams, 1975) and offshore Florida (C13r magnetochron)
(Van Mourik et al., 2001) which are both closer to the
lower Rupelian age estimate of 33.27 Ma (Williams et al.,
2004; 33.44 Ma).
LO Lentinia serrata. The last occurrence of this species at the study site is clearly
observed at ∼48.6 m (159.5 ft), after which only isolated
occurrences are detected. In mid-latitude climatic zones in the Northern
Hemisphere, the age of 33.50 Ma (Williams et al., 2004; 33.68 Ma) has been assigned to this event, whereas in equatorial regions it is
estimated that this taxon persists longer, potentially up to 31.00 Ma
(Williams et al., 2004).
LO Glaphyrocysta semitecta. This species has a fairly continuous occurrence from the base of the
section upwards, becoming particularly common (up to ∼9 %
of the dinocyst assemblage) between ∼115.5 and
∼92.4 m (379 and 303 ft). However, this species is again
common (∼4 % of the dinocyst assemblage) in two upper
intervals (55.8–53.3 m, 183–175 ft; and 41.2–37.5 m, 135–123 ft), which is
unlikely to be a result of reworking and has a last occurrence at
∼37.5 m (123 ft). The LO of G. semitecta has been used in central Italy
to define the base of the Cin dinocyst biozone
(Brinkhuis and Biffi, 1993), of Rupelian age,
which is calibrated with the upper part of P18 planktonic foraminifera
biozones, the upper part of NP21 nannoplankton biozones, and the C13n/C12r
magneto-boundary, which is in agreement with the observations within this
succession.
LO Saturnodinium pansum. This taxon has nearly continuous occurrence throughout the core, only
being absent above ∼20.4 m (67 ft). In sediments recovered
from the central part of the Norwegian Sea (Williams and
Manum, 1999) this extinction event is identified at the Lithostratigraphic
Unit V, corresponding to the upper Rupelian and an age of 30.24 Ma.
LO Enneadocysta arcuata. This species occurs near continuously through almost the whole core,
with highest abundances from ∼105.8 m (347 ft) up to its last
appearance at ∼19.2 m (63 ft). A latest Priabonian age has
been assigned to this bioevent in the Norwegian Sea
(Manum, 1976; Manum et al., 1989),
corresponding to the nannoplankton zone NP21 and the planktonic foraminifera
P17/P18 boundary. However, in the Scotian Shelf and Grand Banks, offshore
southeastern Canada, E. arcuata was an age-diagnostic species for the early Oligocene
D. heterophlycta Assemblage Zone (Williams, 1975). At the Ocean Drilling Program (ODP) Site 1172, in the
East Tasman Plateau (Brinkhuis et al., 2003), an age
estimate of 33.50 Ma (33.68 Ma) has been assigned to the unnamed
Enneadocysta sp. A, a taxon closely related to E. arcuata, which is also consistent with the
placement of this event in North America.
On a regional scale, some distinctive events found at Mossy Grove are very
similar to those detected in five sections from southeastern Mississippi and
southwestern Alabama (Jaramillo and
Oboh-Ikuenobe, 1999). Graphic correlation across five sections (St. Stephens
Quarry, R2089, #1 Young, #1 Wayne, and #1 Ketler) allowed age
determination of the main dinocyst events, including the FO of
Dinopterygium cladoides (33.9 Ma), the Acme of Operculodinium centrocarpum (33.6 Ma), the LO of Dinopterygium cladoides (33.5 Ma), and the LO of
Cribroperidinium tenuitabulatum (33.2 Ma). Although these are not included in the list of the discussed
bioevents above, they are all apparent in the Mossy Grove core, with depths
listed in Table 2, and appear to be consistent in
their age and stratigraphic position with these regional studies.
It is notable that well-preserved specimens of Areosphaeridium diktyoplokum were found at
∼74.1 m (243 ft). The extinction of this taxa has been
associated with the EOB (e.g.
Williams, 1975, 1977; Williams
and Bujak, 1985; Head and Norris, 1989), although there is much discussion
about the age of this event. Early publications reported late Eocene ages
for the event in the Scotian Shelf; Grand Banks (Williams,
1975, 1977); and in Baffin Bay, Canada (Head and Norris,
1989), as well as in the North Sea
(Bujak and Mudge, 1994; Mudge
and Bujak, 1996). However, studies of central Italian successions
(Brinkhuis and Biffi, 1993; Wilpshaar
et al., 1996) place the LO of A. diktyoplokum above the hantkeninid extinction horizon, at
the top of the early Oligocene Adi interval zone, calibrated against lower
planktonic foraminifera P18 zone, mid nannoplankton NP21 zone, and the basal
part of magnetic polarity chronozone C13n. Although previous
(Costa and Manum, 1988; Stover and
Williams, 1988; Williams and Bujak, 1985) and later
(Bujak and Mudge, 1994; Mudge
and Bujak, 1996) surveys have demonstrated the possibility that the alleged
later occurrences in central Italy could be due to reworking processes,
recent dinocyst bio-stratigraphies from the Norwegian–Greenland seas
(Eldrett et al.,
2004; Eldrett and Harding, 2009) confirmed a post-hantkeninid age for the
extinction of A. diktyoplokum, with an age assigned of 33.47 Ma
(Westerhold et al., 2014;
Williams et al., 2004). Although this species is rare in the Mossy Grove
succession, based on correlations to nannofossil biostratigraphy, the
topmost occurrence has an estimated age of 34.12 Ma, which is very close to
the EOB.
Integrated age model
Using these ages, we present a significantly refined age model for the Mossy Grove core (Fig. 5) based
on dinocyst bioevents as discussed in the previous section
(Table 2), integrated with new age constraints from
calcareous nannofossil assemblages as well as existing radiometric dates. We
show previously published foraminiferal-based age control for reference
(Table 1).
Refined age–depth model of the Mossy Grove core. Lithology follows Dockery III et al. (1991) and the timescale follows Westerhold et al. (2014). Best fit lines for the models of Fluegeman et al. (2009) and this work are represented as green and red lines, respectively. Sedimentation rates are also indicated. The shaded area in light red reproduces an error of ∼0.3 Myr, simulating the inclusion of the stated error on radiometric dating of both bentonites. Empty markers refer to bioevents regionally correlated (see Fig. 2).
This new age model can be interpreted to represent sedimentation within two
intervals with distinct rates: between the bottom of the analysed section,
i.e. ∼152.1 m (499 ft; 37.49 Ma) and ∼89.2 m
(292 ft; 34.44 Ma), the sedimentation rate was 2.1 cm kyr-1 and between
∼89.2 m (292 ft; 34.44 Ma) and the top of the analysed
section, i.e. ∼16.8 m (55 ft; 33.05 Ma), the
sedimentation rate was 4.7 cm kyr-1. This demonstrates that the depositional
processes supplied the basin at a relatively constant rate for at least 3 million years, with a significant change only at ∼34.4 Ma.
This new age model also allows us to place the Priabonian–Rupelian boundary
(EOB: 33.89 Ma) at ∼63.4 m (208 ft) and is more consistent
with radiometric dating: the bentonites previously dated at 33.40 and 34.36 Ma here are 33.11 and 34.34 Ma in age, respectively.
Another direct consequence of this model is the reassessment of the age of
some dinocyst bioevents. As already mentioned, the extinction of R. borussica, for
example, is usually correlated with a late Bartonian age in several sites of
the North Sea and the Norwegian–Greenland sea. Meanwhile, middle Priabonian ages
were assigned to this event in the central Danish basin
(Heilmann-Clausen and Van Simaeys, 2005), corroborating similar
dating in the North Atlantic (Costa and Downie, 1979) and
offshore Florida (Van Mourik et al., 2001). The last
occurrence of this taxon at Mossy Grove was observed at ∼36.77 Ma, consistent with a middle Priabonian age. Two other events are
worth mentioning: the earliest appearance of A. alcicornu is similar in age to sites in
the Mediterranean (Brinkhuis and
Biffi, 1993; Wilpshaar et al., 1996), being an event prior to EOB; on the
other hand, the extinction of D. colligerum, here identified in the lower Oligocene, is a
post-EOB event, as suggested by several authors (e.g.
Kothe, 1990; Brinkhuis and Biffi, 1993).
Comparisons with nearby sites in Mississippi and Alabama
(Houben et al., 2018;
Jaramillo and Oboh-Ikuenobe, 1999) (Fig. 3) show a
similar sequence of bioevents as the Mossy Grove section. The following key
events within the Gulf Coast Plain show a high degree of consistency (from
the older to the younger): FO D. cladoides, LO H. semilunifera, FO A. alcicornu, LO D. colligerum, LO L. serrata, and LO C. tenuitabulatum, and they are judged to
be reliable biohorizons for future biozonation of the upper Eocene to lower
Oligocene of the Gulf Coastal Plain.
Conclusions
The Mossy Grove core recorded a rich upper Eocene to lower Oligocene section
with excellently preserved palynomorph content, including 52 genera
and 70 species of dinocysts identified. Furthermore, rich diversity in
other palynological groups was observed, revealing the enormous potential of
the study site for the elaboration of future palynofacies models, useful in
palaeobathymetric studies. High-resolution analysis (∼26 kyr)
of the dinocyst content, integrated with calcareous nannofossil bioevents,
generated a more robust age–depth model than previous models based on
planktonic foraminifera alone. The new age–depth model is also more
consistent with 40Ar/39Ar radiometric dating. On the basis of the
new age model, the cored section spans ∼5 Ma, including a
∼4 Ma of the late Eocene and ∼1 Ma of the
earliest Oligocene, including the critical Eocene–Oligocene transition. The
age model also documents a substantial increase in sedimentation rate from
∼2.1 to ∼4.7 cm kyr-1 at ∼34.4 Ma, indicating that the EOT climatic event caused a major change in the
sedimentary regime.
Data availability
All semi-quantitative data are provided within the body of the article.
List of taxa
Achomosphaera alcicornu (Eisenack 1954) Roger Jack Davey and Williams 1966 (Plate 1, fig. 1)
Illustration of taxa and sample or slide number. (1) Achomosphaera alcicornu (MG55; 16.8 m; U15).
(2) Araneosphaera araneosa (MG479; 146.0 m; C19-1). (3, 4) Areoligera sentosa (MG423; 128.9 m; C7). (5, 6)
Areosphaeridium diktyoplokum (MG243; 74.1 m; C11-1). (7) Charlesdowniea coleothrypta (MG271; 82.6 m; L17-1). (8, 9)
Cleistosphaeridium ancyreum (MG55; 16.8 m; N23-1). (10) Cleistosphaeridium polypetellum (MG499; 152.1 m; D9). (11) Cordosphaeridium cantharellus (MG215; 65.5 m;
C11-1). (12, 13) Cordosphaeridium fibrospinosum (MG383; 116.7 m; D12; MG371; 113.1 m; E18-2). (14, 15)
Cordosphaeridium inodes (MG347; 105.8 m; C15). (16–19) Corrudinium incompositum (MG59; 18.0 m; G37, K11-2). (20)
Cribroperidinium tenuitabulatum (MG363; 110.6 m; C18-2).
Illustration of taxa and sample or slide number. (1) Dapsilidinium pastielsii (MG391; 119.2 m;
D11). (2) Deflandrea heterophlycta (MG111; 33.8 m; C22-3). (3, 4) Deflandrea phosphoritica (MG111; 33.8 m; C10-4). (5, 6)
Dinopterygium cladoides (MG187; 57.0 m; C6). (7) Diphyes colligerum (MG259; 78.9 m; H16). (8, 9) Diphyes ficusoides (MG231; 70.4 m; D20).
(10, 11) Distatodinium biffii (MG71; 21.6 m; E5; MG471; 143.6 m; C18). (12) Distatodinium ellipticum (MG59; 18.0 m;
F24-2). (13) Distatodinium paradoxum (MG103; 31.4 m; D14). (14) Echidinium sp. (MG59; 18.0 m; M42-4). (15, 16)
Enneadocysta arcuata (MG263; 80.2 m; F12). (17) Enneadocysta deconinckii (MG127; 38.7 m; C10). (18, 19) Glaphyrocysta laciniiformis (MG371; 113.1 m;
C8-1). (20) Glaphyrocysta retiintexta (MG427; 130.2 m; E12).
Illustration of taxa and sample or slide number. (1–4) Glaphyrocysta semitecta (MG331; 100.9 m;
C16, C20-4). (5) Hemiplacophora semilunifera (MG471; 143.6 m; C19-1). (6) Heteraulacacysta leptalea (MG487; 148.4 m; D14-4). (7)
Heteraulacacysta porosa (MG111; 33.8 m; C8-2). (8, 9) Homotryblium abbreviatum (MG319; 97.2 m; C10-4). (10–12) Homotryblium floripes (MG55; 16.8 m; K12-4, Q7). (13) Homotryblium oceanicum (MG59; 18.0 m; F8-3). (14) Horologinella pentagonalis (MG319; 97.2 m; O22). (15)
Hystrichokolpoma granulatum (MG63; 19.2 m; B21-4). (16) “Hystrichokolpoma pseudooceanicum” (MG455; 138.7 m). (17) Hystrichokolpoma rigaudiae (MG59; 18.0 m; H24-2).
(18, 19) Hystrichokolpoma salacia (MG299; 91.1 m; E12-4; MG307; 93.6 m; F20-4). (20)
Hystrichosphaeridium tubiferum (MG299; 91.1 m; C6-4).
Illustration of taxa and sample or slide number. (1, 2) Impagidinium paradoxum (MG459; 139.9 m;
D5). (3) Lejeunecysta sp. (MG411; 125.3 m; C10-2). (4, 5) Lentinia serrata (MG59; 18.0 m; N16-4, Q14-3). (6,
7) Lingulodinium “brevispinosum” (MG299; 91.1 m; D12; MG259; 78.9 m; D23-3). (8, 9) Lingulodinium machaerophorum (MG215; 65.5 m; D18;
MG311; 94.8 m; C10-2). (10) Melitasphaeridium pseudorecurvatum (MG199.5; 60.8 m; C14). (11) Minisphaeridium latirictum (MG275; 83.8 m;
K7). (12, 13) Nematosphaeropsis sp. (MG451; 137.5 m; H15-1, J17-3). (14) Operculodinium centrocarpum (MG459; 139.9 m;
D11-1). (15, 16) Operculodinium microtriainum (MG451; 137.5 m; J17-3). (17) Phthanoperidinium comatum (MG323; 98.5 m; E8). (18)
Phthanoperidinium distinctum (MG95; 29.0 m; D8-4). (19) Piladinium columna (MG55; 16.8 m; F20). 20. Rhombodinium sp. endocyst (MG55; 16.8 m;
K11-1).
Illustration of taxa and sample or slide number. (1) Rhombodinium spinula (MG447; 136.3 m;
C21). (2) Rottnestia borussica (MG459; 139.9 m; C13-1). (3, 4) Saturnodinium pansum (MG311; 94.8 m; T8-4; MG411; 125.3
m; C9-4). (5) Senegalinium sp. (MG55; 16.8 m; C15). (6) Sophismatia tenuivirgula (MG55; 16.8 m; E9-3). (7)
Spiniferites bentorii (MG55; 16.8 m; G18-1). (8) Spiniferites mirabilis (MG259; 78.9 m; H11). (9, 10) Spiniferites pseudofurcatus (MG459; 139.9 m;
E7-2). (11) Spiniferites ramosus group (MG59; 18.0 m; E10). (12) Tectatodinium pellitum (MG59; 18.0 m; H9-1). (13, 14)
Thalassiphora spinifera (MG463; 141.1 m; C9; MG363; 110.3 m; C20). (15) Turnhosphaera hypoflata (MG343; 104.6 m; E22). (16)
Wetzeliella articulata (MG319; 97.2 m; N22).
Illustration of taxa and sample or slide number. (1, 2) Svalbardella partimtabulata (MG419; 127.7 m;
D7-3). (3) Thalassiphora fenestrata (MG175; 53.3 m; E21-2). (4, 5) Thalassiphora pelagica (MG499; 152.1 m; C6-2).
Illustration of reworked taxa and sample or slide number. (1)
Achomosphaera ramulifera (MG363; 110.6 m; C20). (2) Cannosphaeropsis sp. (MG167; 50.9 m; C20-1). (3, 4) Chatangiella sp. (MG283;
86.3 m; D10-4; MG271; 82.6 m; L19-3). (5) Cymososphaeridium? phoenix (MG167; 50.9 m; C21-1). (6)
Dinogymnium sibiricum (MG371; 113.1 m; D18). (7) Litosphaeridium siphoniphorum (MG363; 110.6 m; C17-2). (8)
Palaeohystrichophora infusorioides (MG275; 83.8 m; K5-3).
Author contributions
MADLM conceived the study, processed and analysed samples, interpreted results, and wrote and edited the article; GH analysed the samples and participated in interpretation of the article; TDJ collected samples and participated in interpretation, writing, and editing of the article.
Competing interests
The authors declare that they have no conflict of interest.
Acknowledgements
We thank Carlos d'Apolito Júnior (Federal University of Mato
Grosso, Brazil) and Manuel Vieira (Shell, UK), as well as James A. P. Bendle and Kirsty M. Edgar (University of Birmingham, UK) for their
numerous and helpful suggestions, which have greatly improved this work; we thank
Roger Burgess (University of Aberdeen, UK) for the initial pilot study,
revealing the excellent level of preservation of the samples.
We also thank PetroStrat Ltd, particularly Gary Smith, Marcel Polling, Neil Campion, and Paul Cornick, for allowing us access to their facilities during the slide preparation step.
Financial support
This research has been supported by the National Council for Scientific and Technological Development (CNPq, Brazil) (grant no. 206218/2014-1) and the Natural Environment Research Council (NERC, UK) (standard grant NE/P013112/1).
Review statement
This paper was edited by Francesca Sangiorgi and reviewed by Henk Brinkhuis, Kasia K. Sliwinska, and two anonymous referees.
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