Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 320–331
Contents lists available at ScienceDirect
Palaeogeography, Palaeoclimatology, Palaeoecology
j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / p a l a e o
Palaeoclimate perturbations before the Sheinwoodian glaciation: A trigger for
extinctions during the ‘Ireviken Event’
Oliver Lehnert a,b,⁎, Peep Männik c, Michael M. Joachimski a, Mikael Calner d, Jiři Frýda b,e
a
Universität Erlangen, Geozentrum Nordbayern, Schlossgarten 5, D-91054 Erlangen, Germany
Czech Geological Survey, Klárov 3/131, 118 21 Prague 1, Czech Republic
Tallinn University of Technology, Institute of Geology, Ehitajate tee 5, 19086 Tallinn, Estonia
d
GeoBiosphere Science Centre, Lund University, Sölvegatan 12, SE-223 62 Lund, Sweden
e
Faculty of Environmental Sciences, CULS, 165 21 Prague 6, Czech Republic
b
c
a r t i c l e
i n f o
Article history:
Received 8 July 2009
Received in revised form 22 December 2009
Accepted 5 January 2010
Available online 11 January 2010
Keywords:
Palaeoclimate
Glaciation
Oxygen isotopes
Silurian
Baltoscandia
Estonia
a b s t r a c t
Telychian to Sheinwoodian conodonts (Pterospathodus eopennatus ssp. n. 1 Zone through the Ozarkodina
sagitta rhenana Superzone) were studied in detail for their oxygen isotope composition. The Upper Kockelella
ranuliformis and Ozarkodina sagitta rhenana Zones represent the peak interval of the Sheinwoodian positive
δ18Oapatite excursion reflecting the coolest conditions with marine sea–water temperatures decreased by
more than 6 °C in the subtropics. The δ18Oapatite data cover pre- through post-‘Ireviken Event’ strata. The
study includes also material from younger levels up to the late Wenlock Ctenognathodus murchisoni Zone.
However this record is incomplete with one major hiatus reflecting the regression associated with the
Middle Silurian Mulde Event.
The stratigraphic levels of the analyzed conodont samples from Estonia can be correlated with high precision
to the datum points of the ‘Ireviken Event’ strata in the Lower and Upper Visby formations on Gotland,
Sweden. Inferred short-term climatic changes in the upper Telychian part of the studied interval may have
been caused by phases of major volcanic activity along the Caledonian front. Across the ‘Ireviken Event’ we
observe variations in oxygen isotope ratios which allow a direct comparison to the stepwise extinctions and
faunal reorganisations in different groups. Our data suggest that faunal extinctions are connected to time
intervals of warming before the establishment of more stable and cooler conditions during the main
Sheinwoodian glacial. The most severe extinctions and faunal turnovers at datum points 2 and 4 occurred
during the early, warmer interval of the event. After the shift into the Sheinwoodian icehouse in the Lower
Kockelella ranuliformis Zone, reflected by rapidly decreasing sea–water temperatures, faunas restabilized and
reefal communities started to flourish. Sedimentary evidence for the glaciation is discussed for different
palaeolatitudes.
© 2010 Elsevier B.V. All rights reserved.
1. Introduction
For a long time the Silurian was regarded as a period of
environmental stability within a long-lasting greenhouse period.
This view was expressed in the titles of many publications, for
example “Silurian and Devonian reefs: 80 million years of global
greenhouse between two ice ages” by Copper (2002) or “Bathymetric
and isotopic evidence for a short-lived Late Ordovician glaciation in a
greenhouse period” in Brenchley et al. (1994). A wealth of studies on
bioevents and geochemistry have more recently proved the opposite,
that the Silurian is a time period of repeated changes to both fauna
and climate (see summaries by Calner, 2008 and Cramer et al., 2010this issue), starting already in the Late Ordovician. The first evidence
⁎ Corresponding author. Universität Erlangen, Geozentrum Nordbayern, Schlossgarten 5, D-91054 Erlangen, Germany. Tel.: +49 9131 8522632; fax: +49 9131 8529295.
E-mail address: [email protected] (O. Lehnert).
0031-0182/$ – see front matter © 2010 Elsevier B.V. All rights reserved.
doi:10.1016/j.palaeo.2010.01.009
for icehouse conditions and glacial events in this time interval is a
∼2‰ shift in oxygen isotopes of conodont apatite in the uppermost
Sandbian of Laurentia (Buggisch et al., 2010). Another, slightly
younger pre-Hirnantian glaciation is suggested by extensive middle
Katian palaeokarst with preserved palaeorelief in Baltoscandia
(Calner et al., 2010-this issue) and western Laurentia (Keller and
Lehnert, 2010-this issue). The Hirnantian glaciation is well constrained by Gondwanan tillites. It is notable that many of the tillites
were initially believed to be of Silurian age (Schönian 2003). Since the
publication of the classic paper by Brenchley et al. (1994) indicating
that there was only a short-lived terminal Ordovician glaciation, these
earlier papers suggesting the existence of Silurian diamictites have
been overlooked. Dating of tillites by palynomorphs, however, has
conclusively showed that some of the diamictites found in Brazil and
Bolivia are of Silurian age (Grahn and Caputo, 1992; Grahn and Paris,
1992; Grahn et al., 2000; Grahn and Gutiérrez, 2001). Data
summarised in the reviews by Díaz-Martínez and Grahn (2007) and
O. Lehnert et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 320–331
Díaz-Martínez (2007) show a continual latest Ordovician through
early Wenlock record of diamictites in different lithostratigraphic
units of Bolivia (e.g. the Llandovery Cancañiri Formation, the middle
to late Llandovery Zapla Formation, the early Wenlock Kirusillas
Formation) and Grahn and Caputo (1992) dated two Telychian
glaciations in Brazil. Post-Sheinwoodian glacials have been proposed
on the basis of stable isotope records connected with substantial sealevel falls (grey triangles in Fig. 1), e.g. during the Late Wenlock Mulde
Event (Calner and Jeppsson, 2003) and the late Ludlow Lau Event
(Eriksson and Calner, 2008). In the northern hemisphere, at least the
Late Ordovician through early Silurian time interval is characterized
by the collision of Baltica and eastern Avalonia, which together with
Laurentia form the large palaeocontinent Laurussia (Fig. 2A). As
321
indicated by a great abundance of bentonites in contemporaneous
strata in Baltoscandia (e.g. Bergström et al., 1999; Kiipli et al., 2008a,b)
this process was associated with intense volcanism, possibly triggering a change in global climate.
In this paper we present a new early Silurian oxygen isotope
record from the East Baltic Viki core (Fig. 2B). We discuss its
significance for Telychian through Sheinwoodian palaeoclimate and
how such changes may have affected contemporaneous palaeoenvironments and fauna. This time interval is widely recognised for its
extinctions among graptolites and conodonts during the ‘Ireviken
Event’ (sensu Jeppsson, 1998). Locations with published records of
this bioevent and of the carbon isotope excursion in the early
Sheinwoodian are compiled in Fig. 2A. This includes areas such as
Fig. 1. Stratigraphic chart showing the position of the Silurian bioevents (modified from Calner, 2008; original figure from Johnson, 2006). Conodont stratigraphic data are from
Jeppsson (1997a, 2005), Zhang and Barnes (2002), and Männik (2007a,c); oceanic event terminology follows Jeppsson (1990, 1998). For space reasons, the Lansa Secundo Episode
and the Allekvia Primo Episode were excluded from the figure. These are otherwise found between the Boge and Valleviken events. The letters H and A in the right column show the
subdivision of the Gotland succession based on the humid (H) and arid (A) climate periods proposed by Bickert et al. (1997). The early Silurian glaciations (black triangles with a
white G) are based on the presence of tillites in the Amazon and Paraná basins of Brazil and in the Andean basins of Argentina, Bolivia, and Peru (Grahn and Caputo, 1992; Caputo,
1998; see also Díaz-Martínez and Grahn, 2007) as well as on oxygen isotope data (Azmy et al., 1998). Grey triangles indicate levels for which glaciations have been suggested based
on sedimentary data and the stable isotope record (Jeppsson and Calner, 2003; Lehnert et al., 2007).
322
O. Lehnert et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 320–331
Fig. 2. Palaeogeography and locations were the early Silurian IBE and/or the ESCIE has been documented. A. Palaeogeographic reconstruction modified from Cocks and Torsvik
(2002). Black stars indicate areas where the ESCIE together with facies changes and/or biodiversity changes across the IBE has been recorded. White stars show only areas which
include a record of strong faunal and/or facies changes connected with the IBE but no stable isotope data. One star may represent one or several locations in one of the listed areas.
Reports from the Laurentian part of Laurussia include Arctic Canada (AC; Noble et al., 2005), Quebec (QU; Azmy et al., 1998), Alaska (AL; Jeppsson, 1997a), Nevada (NE; Saltzman,
2001), Oklahoma (OK; Saltzman, 2001), Iowa (IO; Cramer and Saltzman, 2005), Ohio (OH; Cramer and Saltzman, 2005), Tennessee (TN; Cramer and Saltzman, 2005), New York and
Ontario (NY; Brand et al., 2006), North West Territories (NWT; Jeppsson, 1997a). The Laurussian record from the Baltoscandian Basin includes Norway (NO; Kaljo et al., 2004),
Sweden (SW; Wenzel and Joachimski, 1996; Munnecke et al., 2003), the East Baltic area (EB; Kaljo et al., 1997, 1998), Podolia (PO; Kaljo et al., 2007), the record from the Avalonian
part of Laurussia treats data from Great Britain (GB; Jeppsson, 1997a; Munnecke et al., 2003; Loydell and Frýda, 2007). There is one study from tropical Gondwana (NSW; New South
Wales, Australia; Talent et al., 1993), and only a few papers treat locations in higher latitudes of peri-Gondwana and Gondwana: Austria (AU; Carnic Alps; Wenzel, 1997), Prague
Basin (CR; Czech Republic; this study), northern Africa (Tunisia, TU; Tunisia; Vecoli et al., 2009), and the northern Greater Caucasus region (CAU; Ruban, 2008). A black triangle
indicates the high latitude position of well dated early Sheinwoodian diamictites in Bolivia (BO; Díaz-Martínez, 2007; Díaz-Martínez and Grahn, 2007). B. Palaeogeographic map of
the Baltoscandian Basin showing the different facies belts of Jaanusson (1976, 1995) and the location of the studied Viki core (1) and the Liva Cliff section (2) on Saaremaa and of the
exposures of the Lower Silurian succession along the northwest coast of Gotland (3) (modified from Nielsen 1995 and Stouge 2004).
Fig. 3. The studied interval in the Viki core section, showing the oxygen isotope record based on conodont phosphate. A detailed conodont zonation of the Viki core succession is
described by (Männik, 2007c). The Valgus and lithuanicus events, described from other core sections in the East Baltic area (Männik 2007b), have been identified by Peep Männik in
the Viki core. Their global impact on faunas is not yet as intensively studied as the three main bioevents of the Silurian (Ireviken, Mulde and Lau). Therefore, the ranges of these
events are indicated by lighter grey bars in contrast to the range of the IBE. The record of some major volcanic ash layers in the core succession is indicated by black rectangles with
the bed numbers shown by Kiipli et al. (2001).
O. Lehnert et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 320–331
323
324
Table 1
Data set from the Viki core—Saarema/Estonia–Telychian through Sheinwoodian.
Collection nos.
Stage (regional stage)
Formation
Depth
(m)
Conodont zonation
(Männik 2007a,c)
PM 1
PM 2
PM 3
PM 4
PM 5
PM 6
PM 7
PM 8
PM 9
PM 10
PM 11
PM 12
PM 13
PM 15
PM 17
OM4-305; M-950
M-951 (1990)
M-953
M-954
M-955 (1990)
M-956 (1990)
M-15 (1982)
M-957
M-959 (1990); M-960
M-961 (1990); M-962
M-963 (1990)
M-14 (1982)
M-964; M-13
C95-69 (1982)
C95-71; OM4-307; C95-72
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Pterospathodus eopennatus Superzone /
Pterospathodus eopennatus Superzone /
Pterospathodus eopennatus Superzone /
Pterospathodus eopennatus Superzone /
Pterospathodus eopennatus Superzone /
Pterospathodus eopennatus Superzone /
Pterospathodus eopennatus Superzone /
Pterospathodus eopennatus Superzone /
Pterospathodus eopennatus Superzone /
Pterospathodus eopennatus Superzone /
Pterospathodus eopennatus Superzone /
Pterospathodus eopennatus Superzone /
Pterospathodus eopennatus Superzone /
Pterospathodus eopennatus Superzone /
Pterospathodus eopennatus Superzone /
PM 18
PM 19
PM 20
PM 22
PM 23
PM 24
PM 25
PM 28
PM 29
PM 30
PM 31
PM 32
PM 35
PM 36
PM 37
PM 38
PM 39
PM 40
PM 41
C95-73; M-10
C95-75; C95-74
M-9; C95-76
C98-2
C95-78; OM4-308
M-8 (1982); C95-79
C95-80
OM4-309
M-6 (1982)
C95-86
M-5 (1982); C95-87
C95-88
M-2 (1982)
M-1 (1982)
M-965
M-966 (1990); OM4-311
M-360 (1984)
M-969
M-970; M971; M362
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
PM 42
PM 43
PM 44
PM 45
PM 47
PM 48
PM 49
PM 53
PM 54
PM 55
PM 58
PM 59
PM 60
PM 61
PM 62
PM 63
PM 64
PM 65
PM 66
M-972
M-974 (1990)
M-975 (1990)
M-977 (1990)
M-981 (1990); M–366
M-982 (1990)
M-367 (1984)
M-374 (1984); OM4-314
M-375
M-376
M-379 (1984); M-380 (184)
M-381 (1984)
M-282
OM4-316; M-383
M-384 (1984); M-385 (1984)
OM4-317; M-386
M-388 (1984); M-387 (1984)
OM4-318; M-389 (1984)
M-390 (1984); M-984 (1990);
OM4-319; M985 (1990)
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
Telychian
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Adavere)
(Jaani)
(Jaani)
(Jaani)
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Velise
Jaani
Jaani
Jaani
Jaani
Jaani
Jaani
Jaani
Jaani
Jaani
184,08–184,18 / 183,90–184,00
183,65–183,80
183,32–183,42
183,17–183,32
183,04–183,17
182,90–183,04
182,75–182,90
182,42–182,54
181,95–182,10 / 181,81–181,91
181,52–181,60 / 181,35–181,45
181,11–181,24
180,60–180,70
179,82–179,95 / 179,55–179,68
177,20–177,30
174,95–175,05 / 174,29–174,40 /
174,20–174,30
173,00–173,15 / 173,55–173,70
172,05–172,15 / 172,50–172,60
171,57–171,70 / 170,95–171,05
169,75–169,85
169,25–169,38 / 169,22–169,37
168,70–168,85 / 168,00–168,10
167,25–167,40
163,95–164,08
163,60–163,75
161,90–162,05
161,40–161,50 / 160,55–160,70
159,80–159,95
155,35–155,50
154,35–154,50
154,15–154,25
153,80–153,97 / 153,71–153,84
153,37–153,50
152,80–152,90
152,14–152,28 / 151,85–152,00 /
151,54–151,64
151,25–151,40
150,80–150,87
150,17–150,28
149,48–149,55
147,83–147,92 / 147,60–147,70
147,15–147,25
146,55–146,65
138,90–139,00 /138,53–138,70
137,95–138,10
136,96–137,10
132,90–133,05 / 131,85–132,00
130,45–130,60
129,20–129,30
127,90–128,00 / 127,45–127,60
125,50–125,65 / 124,10–124,25
123,70–123,80 / 122,70–122,85
121,50–121,60 / 120,60–120,75
119,95–120,06 / 119,60–119,75
118,40.–118,55 / 118,05–118,20 /
117,92–118,05 / 117,05–117,20
δ18O
P. eopennatus
P. eopennatus
P. eopennatus
P. eopennatus
P. eopennatus
P. eopennatus
P. eopennatus
P. eopennatus
P. eopennatus
P. eopennatus
P. eopennatus
P. eopennatus
P. eopennatus
P. eopennatus
P. eopennatus
ssp. n. 1 Zone
ssp. n. 1 Zone
ssp. n. 1 Zone
ssp. n. 1 Zone
ssp. n. 1 Zone
ssp. n. 1 Zone
ssp. n. 1 Zone
ssp. n. 1 Zone
ssp. n. 1 Zone
ssp. n. 1 Zone
ssp. n. 1 Zone
ssp. n. 2 Zone
ssp. n. 2 Zone
ssp. n. 2 Zone
ssp. n. 2 Zone
/ lower Subzone
/ lower Subzone
/ lower Subzone
/ upper Subzone
18,1
18,3
18,4
17,8
17,8
18,2
18,0
17,9
18,6
18,0
17,9
18,4
18,7
18,8
19,1
Pterospathodus eopennatus Superzone / P. eopennatus ssp. n. 2 Zone / upper Subzone
Pterospathodus eopennatus Superzone / P. eopennatus ssp. n. 2 Zone / upper Subzone
Pterospathodus eopennatus Superzone / P. eopennatus ssp. n. 2 Zone / upper Subzone
Pterospathodus celloni Superzone / P. amorphognathoides angulatus Zone, Lower Subzone
Pterospathodus celloni Superzone / P. amorphognathoides angulatus Zone, Lower Subzone
Pterospathodus celloni Superzone / P. amorphognathoides angulatus Zone, Lower Subzone
Pterospathodus celloni Superzone / P. amorphognathoides angulatus Zone, Lower Subzone
Pterospathodus celloni Superzone / P. amorphognathoides angulatus Zone, Lower Subzone
Pterospathodus celloni Superzone / P. amorphognathoides angulatus Zone, Lower Subzone
Pterospathodus celloni Superzone / P. amorphognathoides angulatus Zone, Upper Subzone
Pterospathodus celloni Superzone / P. amorphognathoides angulatus Zone, Upper Subzone
Pterospathodus celloni Superzone / P. amorphognathoides angulatus Zone, Upper Subzone
Pterospathodus celloni Superzone / P. amorphognathoides angulatus Zone, Upper Subzone
Pterospathodus celloni Superzone / P. amorphognathoides angulatus Zone, Upper Subzone
Pterospathodus celloni Superzone / P. amorphognathoides angulatus Zone, Upper Subzone
Pterospathodus celloni Superzone / P. amorphognathoides lennarti Zone
Pterospathodus celloni Superzone / P. amorphognathoides lennarti Zone
Pterospathodus celloni Superzone / P. amorphognathoides lennarti Zone
Pterospathodus celloni Superzone / P. amorphognathoides lithunicus Zone
18,4
18,2
18,2
18,2
18,6
18,2
18,2
17,5
18,1
17,9
18,5
18,4
18,1
17,9
18,5
18,9
19,7
18,2
19,2
Pterospathodus celloni Superzone / P. amorphognathoides lithuanicus
Pterospathodus celloni Superzone / P. amorphognathoides lithuanicus
Pterospathodus celloni Superzone / P. amorphognathoides lithuanicus
Pterospathodus celloni Superzone / P. amorphognathoides lithuanicus
Pterospathodus celloni Superzone / P. amorphognathoides lithuanicus
P. amorphognathoides amorphognathoides Zone / Lower Subzone
P. amorphognathoides amorphognathoides Zone / Lower Subzone
P. amorphognathoides amorphognathoides Zone / Lower Subzone
P. amorphognathoides amorphognathoides Zone / Lower Subzone
P. amorphognathoides amorphognathoides Zone / Lower Subzone
P. amorphognathoides amorphognathoides Zone / Upper Subzone
P. amorphognathoides amorphognathoides Zone / Upper Subzone
P. amorphognathoides amorphognathoides Zone / Upper Subzone
P. amorphognathoides amorphognathoides Zone / Upper Subzone
P. amorphognathoides amorphognathoides Zone / Upper Subzone
P. amorphognathoides amorphognathoides Zone / Upper Subzone
P. amorphognathoides amorphognathoides Zone / Upper Subzone
P. amorphognathoides amorphognathoides Zone / Upper Subzone
P. amorphognathoides amorphognathoides Zone / Upper Subzone
18,7
17,9
19,3
18,8
18,1
18,3
18,5
18,7
18,0
18,7
18,5
18,6
18,7
18,1
18,3
18,4
18,7
18,8
18,2
Zone
Zone
Zone
Zone
Zone
O. Lehnert et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 320–331
Lab no.
PM 67
M-986 (1990); M-392 (1984);
M-987
Telychian (Jaani)
Jaani
PM 68
M-989 (1990);
(1984)
M-991 (1990);
M-992 (1990);
M-394 (1984);
M-994 (1990)
M-990; M-391
Telychian (Jaani)
Jaani
M 393 (1984);
M993 (1990)
OM4-322;
Telychian (Jaani)
Jaani
Telychian (Jaani)
Jaani
PM 69
PM 70
116,70–116,85 / 116,30–116,45 /
116,10–116,20 / 115,95–116,00 /
115,85–115,95
115,60–115,75 / 115,30–115,45 /
115,10–115,25
115.10-115,23 / 114,90–115,05 /
114,65–114,80 / 114,30–114,45
114,10–114,20 / 113,95–114,05 /
113,80–113,95
P. amorphognathoides amorphognathoides Zone / Upper Subzone
18,7
P. amorphognathoides amorphognathoides Zone / Upper Subzone
18,9
P. amorphognathoides amorphognathoides Zone / Upper Subzone
19,1
P. amorphognathoides amorphognathoides Zone / Upper Subzone
18,5
Datum 1—Ireviken
PM 71
PM 72
M-995 (1990); M-395 (a) (1984)
M-996 (1990)
Telychian (Jaani)
Telychian (Jaani)
Jaani
Jaani
113,75–113,90 / 113,30–113,45
113,25–113,40
Pseudoonetodus bicornis Superzone / Lower Pseudoonetodus bicornis Zone
Pseudoonetodus bicornis Superzone / Lower Pseudoonetodus bicornis Zone
PM 73
M-997; OM4-323
Sheinwoodian (Jaani)
Jaani
113,05–133,20 / 112,95–113,05
Pseudoonetodus bicornis Superzone / Upper Pseudoonetodus bicornis Zone
PM 74
M-1000; M-999; M-998 (1990)
Sheinwoodian (Jaani)
112,80–112,95 / 112,10–112,25 /
112,35–112,45
Pterospathodus p. procerus Superzone / Lower Pterospathodus
p. procerus Zone
PM 75
M-1001; M-395 (b) (1984);
M-1002
Sheinwoodian (Jaani)
18,4
18,9
Datum 2—Ireviken
Datum 3—Ireviken
18,3
Datum 4? or additional datum—Ireviken
Jaani
111,85–112,00 / 111,65–111,80 /
111,45–111,55
Pterospathodus p. procerus Superzone / Upper Pterospathodus
p. procerus Zone
18,4
Datum 6—Ireviken
PM 76
M-1004; M-396 (1984);
M-1005; M-1006
Sheinwoodian (Jaani)
Jaani
110,85–111,00 / 110,25–110,40 /
110,62–110,72 / 110,50–110,62
Kockelella ranuliformis Superzone / Lower Kockelella ranuliformis Zone
18,2
Datum 7—Ireviken
PM 77
PM 79
M-195 (1984);
M-198 (1984)
M-199 (1984);
M-200 (1984);
M-202 (1984);
PM 80
PM 81
PM 82
V. Nestor; C01-100
C01-101; V. Nestor
C01-104; V. Nestor; C01-105
Sheinwoodian (Jaagarahu)
Sheinwoodian (Jaagarahu)
Sheinwoodian (Jaagarahu)
Jaagarahu
Jaagarahu
Jaagarahu
PM 83
PM 84
C01-107; V. Nestor
C01–108; V. Nestor; C01-109
Sheinwoodian (Jaagarahu)
Sheinwoodian (Jaagarahu)
Jaagarahu
Jaagarahu
PM
PM
PM
PM
85
86
87
88
V. Nestor
C01-118
C01-121
V. Nestor; C04-70; C04-71
Sheinwoodian (Jaagarahu)
Sheinwoodian (Jaagarahu)
Sheinwoodian (Jaagarahu)
Sheinwoodian (Jaagarahu)
Jaagarahu
Jaagarahu
Jaagarahu
Jaagarahu
PM
PM
PM
PM
PM
PM
PM
PM
89
90
91
92
93
94
95
96
C04-73
V. Nestor
V. Nestor
V. Nestor
V. Nestor
C04–83
C04-87
C04-91
Sheinwoodian (Jaagarahu)
Sheinwoodian (Jaagarahu)
Sheinwoodian (Jaagarahu)
Sheinwoodian (Jaagarahu)
Homerian (Rootsiküla)
Homerian (Rootsiküla)
Homerian (Rootsiküla)
Homerian (Rootsiküla)
Jaagarahu
Jaagarahu
Jaagarahu
Jaagarahu
Rootsiküla
Rootsiküla
Rootsiküla
Rootsiküla
PM 78
OM4-329;
Sheinwoodian (Jaani)
Jaani
OM4-331;
M-201 (1984)
OM4-332; C01-90
Sheinwoodian (Jaani)
Jaani
Sheinwoodian (Jaani)
Jaani
105,45–105,55 / 104,75–104,85 /
102,20–102,35
101,35–101,50 / 100,88–100,99 /
100,40–100,60 / 99,75–99,95
99,00-99,20 / 98,82–98,97 /
98,25–98,35
88,40–88,60 / 88,10–88,20
87,40–87,55 / 86,70–86,80
84,20–84,35 / 84,00–84,10 /
83,35–83,50
81,65–81,75 / 81,50–81,60
80,50–80,65 / 79,60–79,70 /
79,25–79,45
70,30–70,40
68,45–60
64,45–60
62,70–62,90 / 61,60–61,75 /
59,60–59,75
56,45–56,55
56,20–56,40
50,00–50,20 / 48,60–48,70
45,00–45,10 / 44,00–44,20
30,50–30,70
27,65–27,80
21,85-22,00
at about 14,50
Kockelella ranuliformis Superzone / Upper Kockelella ranuliformis Zone
19,9
Kockelella ranuliformis Superzone / Upper Kockelella ranuliformis Zone
19,8
Kockelella ranuliformis Superzone / Upper Kockelella ranuliformis Zone
19,7
Ozarkodina sagitta rhenana Superzone
Ozarkodina sagitta rhenana Superzone
Ozarkodina sagitta rhenana Superzone
19,3
19,6
19,7
Ozarkodina sagitta rhenana Superzone
Ozarkodina sagitta rhenana Superzone
19,6
19,8
Upper Kockelella
Upper Kockelella
Upper Kockelella
Upper Kockelella
Superzone
Superzone
Superzone
Superzone?
19,6
19,4
19,7
19,6
Upper Kockelella walliseri Superzone?
Upper Kockelella walliseri Superzone?
Upper Kockelella walliseri Superzone?
Upper Kockelella walliseri Superzone?
Ctenognathodus murchisoni Zone
Ctenognathodus murchisoni Zone
Ctenognathodus murchisoni Zone
Ctenognathodus murchisoni Zone
19,8
19,5
19,3
19,0
19,3
19,4
19,6
19,7
walliseri
walliseri
walliseri
walliseri
O. Lehnert et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 320–331
18,7
325
326
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northern Africa (Vecoli et al., 2009) and the Greater Caucasus region
(Ruban 2008) where these events have been documented more
recently. Although they may be closely related, we want to stress the
importance of separating the biotic events from the carbon and
oxygen isotopic events since confusion may arise if these perturbations are assembled under the same name. These perturbations
should be clearly separated, simply because they span different time
slices. With regard to the ‘Ireviken Event’, the biotic event (which is
the original basis for its definition) spans the Lower P. bicornis
through the top of the Lower K. ranuliformis Zone. Hereafter, we refer
to it as the Ireviken BioEvent (IBE). The associated isotope excursions
start considerably later, and there is even a clear time offset between
the start of the positive shifts in δ13C (Lower P. procerus zone) and the
shift in δ18O (Lower K. ranuliformis zone; Fig. 4). In this paper we refer
to them as the Early Sheinwoodian Carbon Isotope Excursion (ESCIE)
and Sheinwoodian Oxygen Isotope Excursion (SOIE), respectively.
from Norway (Kaljo et al., 2004), followed by values at about 5.5‰ in
New York State and Ontario (Brand et al., 2006), about 5‰ in Sweden
(Gotland; Munnecke et al., 2003), about 4‰ in Oklahoma, and about 3%
in Nevada (Saltzman, 2001). δ13Corg data from the North African “hot
shales” of Tunisia display a positive shift of about 3‰ (Vecoli et al., 2009)
and the δ13Corg dataset from Wales shows a shift of more than 3‰
(Loydell and Frýda, 2007). Within the graptolithic, deeper water facies
of the Baltoscandian Basin, the δ13C excursion starts in the upper Cyrtograptus murchisoni graptolite Zone and ends in the Middle Wenlock
(Loydell, 2007). In the carbonate platform facies of Gotland, the shift to
higher δ13C values starts just below datum point 4 of the IBE within the
corresponding Lower Pterospathodus procerus conodont Zone. The onset
of the shift in δ13C at this level has also been reported from other
palaeocontinents (e.g., Talent et al., 1993; Cramer and Saltzman, 2005,
Cramer et al., 2010). It is notable that the peak interval of the excursion
shows a significant offset to the time span of the IBE. For a recent
summary of studies on the ESCIE see also Calner (2008).
1.1. The Ireviken BioEvent (IBE)
1.3. The Sheinwoodian Oxygen Isotope Excursion (SOIE)
A great number of marine faunal groups were affected during this
early Sheinwoodian (Early Wenlock) bioevent (Jeppsson and Männik,
1993; Aldridge et al., 1993; Jeppsson et al., 1995; Jeppsson, 1997b, 1998;
summarised by Calner, 2008). By definition the IBE starts at the base of the
Lower Pseudooneotodus bicornis Zone and ranges through the Lower
Kockelella ranuliformis Zone. It has been calculated that the whole IBE,
including eleven datum points (extinction levels), lasted not more than
0.2 Ma, and that the first datum points (including the most severe
extinction events at datum 2 and 4) fall within the first 0.1 Ma of this time
interval (Jeppsson, 1997a; Cramer et al., 2010). There is a severe decrease
in the diversity of conodont species across the event and 48 out of 60
pandemic species became extinct (Aldridge et al., 1993; Jeppsson, 1997a).
A dramatic extinction is observed in the graptolites (Cyrtograptus
murchisoni Event) of which the global fauna was reduced by 20% (Melchin
et al., 1998; Noble et al., 2005). Among the primary producers, acritarchs
went through a significant reorganisation with an extinction of more than
40 species and an origination of more than 50 species (Gelsthorpe, 2004).
Based on the Gotland succession, the acritarch extinctions occurred
throughout the stratigraphic range of this bioevent. It is, however,
questionable if the bulk of the acritarch extinctions (N80%) in the topmost
four meters of the Visby Formation (Gelsthorpe, 2004) is not just an effect
of changing environmental conditions, as indicated by a shift from
subtidal marls to reef and grainstone facies in the studied sections. Drastic
changes in the composition of faunal assemblages together with a major
drop in diversity are also observed within the chitinozoans (e.g., Nestor et
al., 2002; Hints et al., 2006). Among benthic communities, there is at least
a clear regional diversity decrease in trilobites, recording a loss of ∼50%
(Ramsköld, 1985). Biodiversity changes are also observed within the
polychaete faunas, which were reduced by more than 20% (Eriksson,
2006). There are reorganisations in brachiopod associations together with
a marked decrease of shell accumulations, but only minor extinctions are
observed within this group with respect to the number of taxa being
eliminated (Kaljo et al., 1995). Valentine et al. (2003) reported a marked
change in lingulid brachiopod faunas during the Ireviken Event. The
compilation of palaeontological data (e.g., Robison, 1965; Obut et al.,
1988) from the northern Greater Caucasus region by Ruban (2008)
reveals that the IBE was also a pronounced extinction event in this part of
peri-Gondwana. In the Greater Caucasus area, regarded as an independent
Greater Caucasus terrane in the Hun Superterrane assemblage (Ruban et
al., 2007), the IBE is reflected by the most prominent biotic crisis in the
Silurian graptolitic succession (Ruban, 2008).
1.2. The Early Sheinwoodian Carbon Isotope Excursion (ESCIE)
Based on an increasing number of publications treating the ESCIE,
there is no doubt about the global character of this change in the carbon
cycle. The highest δ13C peak values, reaching 6.6‰, have been reported
In this paper, we present the first detailed oxygen isotope data set
from the Telychian through the early Sheinwoodian. The δ18O record
presented in this paper starts well below the IBE, spans the IBE and
includes also post-IBE strata. For the lower Silurian, only δ18O data
derived from brachiopods of Gotland (Bickert et al., 1997; Samtleben
et al., 2000; Munnecke et al., 2003) have been available across the
SOIE. However, on Gotland the δ18O record starts in the Upper P.
bicornis Zone, i.e. within the range of the IBE and provides no baseline
data before this oxygen isotope event. The SOIE has not yet been
documented from any other area.
2. Sampling and methods
Our palaeoclimate study is based on oxygen isotopes from 87
conodont samples from the Viki core and the Liiva Cliff section of the
Saarema Island (western Estonia; Fig. 2B). A conodont color alteration
index of 1 implies only a very minor thermal alteration to the studied
succession. The lithologic units in the Viki core do not display any
major sedimentological changes and the deep water marls show only
minor variations in carbonate content. Across the IBE interval there is
only a gradual change, from more argillaceous carbonates in the lower
part to more calcareous rocks in the upper part of the interval, but
without sharp lithological changes or evidence for erosion or any
hiatus. Unfortunately, there are no detailed sedimentological and
microfacies studies from this succession, but the general facies
suggests that it is a deeper water succession across the IBE.
The detailed Telychian conodont zonation, based on the Pterospathodus lineage, by Männik (1998, 2007a) provide a means to
correlate the sampled succession with the extremely finely subdivided Gotland succession—which has a zonation that has been
successively refined since the early papers by Martinsson (1967),
Fåhraeus (1969) and Jeppsson (1975, 1979), and not least over the
last decade (e.g. Jeppsson et al., 1994; Jeppsson, 1997b, 2005, 2008;
Jeppsson and Aldridge, 2000; Jeppsson et al., 2006). The stratigraphic
position of our samples is plotted in Fig. 3, and the δ18O values are
summarised in Tables 1 (Viki core) and 2 (Liiva cliff section). For the
Viki core (Ø 12 cm) it was commonly difficult to acquire enough
conodont material for isotope analysis and, as indicated in Table 1,
many samples were therefore lumped. The data points are plotted in
the middle part of the stratigraphic range of the combined samples.
However, if someone would have sampled for such an analysis in a
conventional way directly from the core, longer parts of the core
would have been taken as conodont samples from the beginning. The
samples from the Viki core represented in the conodont collections of
Peep Männik at Tallinn University usually represent only 10 cm of
depth in the cutted core (by using only half the core diameter). In
O. Lehnert et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 320–331
327
Table 2
Data across datum point 7 of the Ireviken Event in the Liiva Cliff section (Saarema, Estonia).
Lab no.
Collection nos.
Stage
Formation
Depth
(m)
Conodont zonation
(Männik 2007a,c)
δ18O
PM 97
PM 98
C02-161
C02–163
Sheinwoodian (Jaani)
Sheinwoodian (Jaani)
Jaani
Jaani
4,12 m below top
Kockelella ranuliformis Superzone / Lower Kockelella ranuliformis Zone
Kockelella ranuliformis Superzone / Lower Kockelella ranuliformis Zone
19,0
19,7
PM 99
PM 100
C02-162
C04-67
Sheinwoodian (Jaani)
Sheinwoodian (Jaani)
Jaani
Jaani
Datum 7—Ireviken
most cases, such a sample size would not have been sufficient to
recover enough conodont phosphate for an analysis. In our analyses
we have had an unusually good control on zonal boundaries and
datum points. In the case of lumping of material we made sure that
such boundaries never have been crossed (for details see Table 1).
Stratigraphically, the studied samples span the interval from the
upper Subzone of the Pterospathodus eopennatus ssp. n. 1 Zone
through the Ozarkodina sagitta rhenana Superzone and cover prethrough post-IBE strata.
The conodont elements (0.5 to 1 mg per sample) were dissolved in
nitric acid and chemically converted to Ag3PO4 using the method
described by Joachimski et al. (2009). The oxygen isotope ratios
were measured on CO (carbon monoxide) using a High Temperature
Conversion Elemental Analyser (TC15 EA) connected online to a
ThermoFinnigan Delta Plus mass spectrometer. All δ18O values are
reported in per mil relative to V-SMOW (Vienna Standard Mean Ocean
Water). Accuracy and reproducibility (1σ = ±0.2‰) were monitored by
multiple analyses of NBS120c and several phosphate reference samples
(for details, see Joachimski et al., 2009). Palaeotemperature calculations
are based on the equation published by Kolodny et al. (1983).
3. Results
3.1. The δ18O record during SOIE
For the entire studied interval, including the IBE, we observe
distinct and substantial variations in oxygen isotope ratios of
conodont apatite. In the lowermost samples from the topmost
Rumba Formation and the basal Velise Formation, δ18O values range
between 18 and 18.5‰ (Fig. 3). There is a pronounced increase to
about 19.1‰ in the uppermost Pterospathodus eopennatus Superzone
followed by an interval of low values, showing a minimum of 17.7‰,
within the Pterospathodus amorphognathoides angulatus Zone of the
Pterospathodus celloni Superzone. Values rise again in the upper part
of the P. a. angulatus Zone to a maximum of 19.2‰. They then scatter
between high values around 19.2‰ and values below 18‰ in the
interval spanning the top of P. a. angulatus Zone through the Lower P.
a. amorphognathoides Zone. In the remainder of this zone and up to
datum point 1 of the IBE, δ18O values range from 18.0 to 18.8‰
(Fig. 3). Across the IBE including the Lower Pseudooneotodus bicornis
through the Lower Kockelella ranuliformis Zone, a few short-term
positive and negative shifts in δ18O are observed. A pronounced and
very rapid positive shift of the longer-term SOIE starts within the
Lower K. ranuliformis Zone. In the uppermost part of that zone the first
higher δ18O value of 18.2‰ is recorded, followed by an interval with
δ18O values between 19.3 and 19.8‰ continuing into the Upper K.
walliseri Superzone. There is a hiatus between the O. sagitta rhenana
Superzone and the Upper K. walliseri Superzone (Fig. 3). This hiatus
corresponds well with an interval of extremely shallow-water facies,
including several unconformities, on Gotland (the Tofta through
Hangvar interval; see Jeppsson, 2008, and own observations). In
contrast to the previously published δ18O record based on brachiopods from Gotland (Bickert et al., 1997; Samtleben et al., 2000;
Munnecke et al., 2003; compiled by Calner et al., 2004), where δ18O
values decrease in the Lower K. walliseri Superzone, data from the
Kockelella ranuliformis Superzone / Upper Kockelella ranuliformis Zone
Kockelella ranuliformis Superzone / Upper Kockelella ranuliformis Zone
19,5
19,9
δ18O record based on conodont material from the Viki core do not
decrease before the Upper K. walliseri Zone. The succession including
the Upper K. walliseri Superzone up to the hiatus just below the Ctenognathodus murchisoni Zone presumably correlates with the succession spanning the Boge Event to the Allekvia Primo Episode on
Gotland and shows slightly lower δ18O values. The Mulde Event is not
recorded in the Viki core due to a hiatus below the Ct. murchisoni
Zone, but values in the latter zone are still above 19.5 ‰.
3.2. Interpretation of the SOIE and its relation to the IBE
Based on our data set, the faunal extinctions during the IBE appear to
occur in warmer periods expressed by lower δ18O values or within
intervals with short-term and repeated climatic variations (Figs. 3 and 4).
The δ18O data from Estonia are correlated to the zonation and event
stratigraphy on Gotland. Before datum point 1, cooler sea–water
temperatures in the latest Telychian are indicated by higher δ18O
values (Fig. 4). After some cooling in the upper P. amorphognatoides
amorphognathoides Zone we see a short-term warming in the uppermost
part of this conodont zone just before datum point 1 (decrease in δ18O).
The first step of extinction at datum point 1 affected mainly the
hemipelagic groups. The warm interval continues into the Lower P. bicornis Zone, but sea–water temperature decreases again just before datum
point 2. Afterwards we observe a continuous warming trend into the
Lower P. procerus Zone (a level above datum point 3). This continuous
warming coincides with some of the most severe extinctions of the IBE
strongly affecting conodonts and trilobites. Then, the warm sea–water
temperatures persist into the lowermost part of the Lower K. ranuliformis
Zone (a level above datum point 6). At datum point 4 (Phaulactis layer on
Gotland), which falls into that warmer interval, extinctions are observed
in inner shelf groups such as corals, ostracods and brachiopods. Below
datum point 7, δ18O values start to increase (by ∼0.9‰) indicating that
the cooling associated with the Sheinwoodian glaciation started in the
upper part of the IBE, with an initial temperature drop of about 4 °C.
Accordingly, the main glacial post-dates the severe extinctions during
datum points 2 and 4. Using earlier isotope studies based on brachiopods
from Gotland (Samtleben et al., 2000; Munnecke et al., 2003) the O. s.
rhenana Zone represents the peak interval of the SOIE. Our data confirm
high values in this interval although we observe maximum values in the
preceding Upper K. ranuliformis Zone. The hiatus at the boundary between
the O. s. rhenana Zone and the Upper K. walliseri Superzone makes it
difficult to judge if this time interval represents maximum values.
Nevertheless, the coolest climatic conditions during the Sheinwoodian
glaciation occurred in this time interval, shortly but clearly after the biotic
extinctions. δ18O values remain high after datum point 7 (with only slight
variations in the range of 0.5 ‰) indicating that cold conditions were
sustained after the IBE. The high δ18O values in the Upper K. walliseri Zone
(Fig. 3) suggest that the Sheinwoodian glaciation presumably persisted up
to this level well above the IBE.
4. Discussion of the pre-IBE interval
The lower part of the studied succession represents an interval
of major volcanic activity in Baltoscandia (Kiipli et al., 1997, 2001;
Kiipli and Kallaste, 2002, 2006; Kiipli et al., 2006, 2007; Kallaste and
328
O. Lehnert et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 320–331
Fig. 4. Stratigraphic range of the IBE (grey shading) showing the position of datum points in the sequence stratigraphy column, and changes in conodont diversity by species
numbers (grey—globally, black—Gotland; based on Jeppsson, 1997b). The figure is modified from the compilation by Calner (2008, fig. 6) based on Munnecke et al. (2003). The
graptolite zonation is based on the data of Loydell (1998), Loydell et al. (2003) and Männik (2007a). The Phaulactis layer (monospecific rugose coral biostrome that is an excellent
marker horizon) represents a maximum flooding interval during a warm period prior to the Sheinwoodian glaciation. Thick dashed vertical line to the right of the conodont diversity
plot indicates the glaciation as proposed by Brand et al. (2006). Black solid bar with “G” to the left of the conodont diversity plot refers to the glacial period suggested in this paper
(starting in the K. ranuliformis Zone and ranging up into the K. walliseri Superzone). References to biological groups indicated in the column with the δ13C data are given in the text.
Kiipli, 2006; Kiipli, 2008). The isotopically ‘noisy’ intervals in the
Telychian seem to be related to times of major volcanic eruptions,
which strongly exceed the intensity of any eruption in the historic
record (Fig. 3). Kiipli et al. (2008b) reported about 80 K-bentonite
beds in the Telychian of Baltoscandia. In an early study, Kiipli et al.
(2001) numbered ∼ 30 important ash beds including the “O” bed,
which is the widely distributed Osmundsberget K-bentonite of
Bergström et al. (1999). Some of these layers represent major
eruptions and their stratigraphic positions in the Viki core (Kiipli
et al., 2001) are indicated in Fig. 3. This intense upper Llandovery
volcanism is related to the successive closure of the Tornquist Sea
and the collision between Baltica and eastern Avalonia (e.g. Trench
and Torsvik, 1992).
Fig. 3 displays that the volcanic activity clusters in two time
intervals (P. eopennatus Superzone and upper P. celloni Superzone).
On a general scale these are characterized by short-term shifts in the
oxygen isotope record, interpreted as short climatic perturbations.
These superimposed on a comparably warm period before the strong
shift into stable icehouse conditions and the Sheinwoodian glaciation.
The earlier volcanic eruptions are interpreted to be responsible for
the cold conditions within the P. eopennatus Superzone. This cooling
is followed by a brief time interval of relative warming in the
uppermost part of this superzone before the δ18O values again imply
renewed cooling in the upper part of the overlying P. celloni
Superzone. This latter time interval is characterized by rapid positive
18
δ O shifts in the range of about 1‰ reflecting substantial drops in
sea–water temperatures in the Baltoscandian Basin of up to 4 °C.
Overall cooler conditions follow this volcanically active period during
the P. a. amorphognathoides zone for which Cramer and Saltzman
(2005) introduced an Upper Telychian glaciation. However, since this
particular time interval is not different in its isotopic development
from the investigated, underlying strata, it is not treated as a separate
glacial herein. The δ18O values in their glacial interval are far below
the values in dataset from the interval interpreted as the time of the
Sheinwoodian glaciation in this paper (approx. 1‰ difference; Fig. 3).
The observed shifts in the time before the Upper K. ranuliformis Zone
may represent an interval of rapid changes between short-term
waning and waxing of ice volume stored on Gondwana and their
deglaciation before the major and long-lived glacial period during the
Sheinwoodian.
5. Sedimentary evidence for a Sheinwoodian glaciation
5.1. Sea-level changes
In contrast to a series of well-established events among graptolites
(e.g., Melchin et al., 1998), detailed information of early Silurian
bioevents in carbonate platform environments are lacking before the
IBE. Detailed research on non-graptolitic faunas across the IBE has
been carried out mainly on Gotland where event stratigraphic
research has been intense since the publication of the oceanic event
model by Jeppsson (1990) and the subsequent documentation of
stepwise faunal extinctions (e.g. Aldridge et al., 1993; Jeppsson and
Aldridge, 2000; Jeppsson and Calner, 2003; Calner, 2005a,b). Several
other areas have also been well-studied with respect to faunal and
isotopic changes across the IBE and ESCIE (see compilations in Cramer
et al., 2010). However, it turns out that little information is available
concerning contemporaneous facies shifts. In the following, we
discuss selected areas where the sedimentary record shows evidence
for sea-level changes and a major sea-level drop related to the time
interval for which an Sheinwoodian glaciation is proposed (Figs. 3 and
4).
There are several reconstructions of sea-level changes through the
studied interval available in the recent literature. The Silurian global
sea-level curve of Johnson (2006; included in Fig. 1) and other
published curves treating pre- through post-event strata (Loydell,
1998; Cramer and Saltzman, 2005) are highly distinct from each other
and partly also in contradiction to observations in Baltoscandia.
The global sea-level changes as proposed by Loydell (1998, and
slightly expanded 2007) fit best with the observed sedimentary facies
on Gotland and our stable isotope record. According to this curve
(Loydell, 2007, Fig.1), there is a distinct regression from a high sea-level
starting in the Cyrtograptus murchisoni Zone and continuing into the
basal Monograptus riccartonensis Zone. This is followed by a small rise in
O. Lehnert et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 320–331
sea-level during the M. riccartonensis Zone followed by a short-lived
regression at the top of that zone. After a new small rise just above the
M. riccartonensis Zone a more prolonged regression continues into the
mid-Wenlock. Based on the sedimentary record on Gotland we
interpret the Phaulactis layer, which is on the top of the highly
argillaceous Lower Visby Formation, as part of a maximum flooding
interval. The overlying Upper Visby Formation includes successively
more packstone and grainstone beds and is clearly regressive.
The slightly younger riccartonensis Event correlates with the K.
ranuliformis Superzone and falls in the early time of the proposed
glaciation in this paper (Fig. 4; post-IBE). This corresponds to the
Högklint Formation on Gotland, which represents a reef platform
truncated on the top by a distinct unconformity. The unconformity is
related to a hiatus in the lowermost O. s. rhenana Zone (Jeppsson,
1998). The development of a reef platform during the main glacial is
partly in conflict with the common growth of reefs during
transgression and highstand of sea-level. Further studies are needed
to test if this is a reef tract that developed during a falling stage of sealevel or during a lowstand situation. The Högklint Formation and
especially the overlying Tofta Formation is exceedingly rich in
microbial carbonates. This microbial mass occurrence correlates well
with the main glacial. It is possible, but needs to be tested, that the
small sea-level rise in the M. riccartonensis zone followed by
regression at the top of this graptolite Zone as proposed by Loydell
(1998) correlates to the short-term slight warming at the end of the
Upper K. ranuliformis Zone (lower δ18O values) followed by colder
temperatures during the O. s. rhenana Zone.
5.2. Evidence of the Sheinwoodian glaciation in high latitude
peri-Gondwana
Of major interest in context of the timing of Sheinwoodian
glaciation is the formation, erosion and re-deposition of the
shallow-water Sacta Limestone Member (Kurasilas Formation) in
Bolivia, located in high latitudes of western peri-Gondwana at this
time. The conodont association in the Sacta Limestone (Merino, 1991)
can be correlated with the O. s. rhenana Superzone. The characteristics
of these cool-water carbonates (microfacies, faunal and non-skeletal
grain content) point to deposition in a shallow shelf setting (DíazMartínez, 2007), not displaying climate warming but the start of the
regression and sedimentation in a shallow environment. In the
Cordillera del Tunari (central Eastern Cordillera of Bolivia) slabs of the
Sacta Limestone Member are resedimented within diamictites (DíazMartínez, 2007). Their transport took place after the basal Sheinwoodian (after deposition and lithification). The diamictites in the
Kurasilas Formation clearly postdate the deposition of the Sacta
Member and thus may be correlated to the late O. s. rhenana
Superzone, the K. walliseri Superzone or may even be of a younger
age. They provide evidence for an Sheinwoodian glaciation rather
than a late Telychian glacial as proposed by Cramer and Saltzman
(2005) (compare Fig. 3). We suggest that these sediments formed
during the period when oxygen isotopes reflect the coldest surface sea
water temperatures during the Wenlock (O. sagitta rhenana Zone
through K. walliseri Superzone).
The Lower Silurian succession in the Prague Basin of northern periGondwana is exclusively composed by fine-grained siliciclastics. The
first lenses and beds of micritic and bioclastic limestones with
hitherto undescribed trilobites, occur in the M. riccartonensis Zone.
The same signal of presumably warmer conditions in the basin was
documented in the vicinity of Prague at this stratigraphical level by
Bouček (1937), Řeporyje area) and Havlíček and Štorch (1990, Malá
Chuchle). It is reflected by this first occurrence of limestone lenses of
the C. murchisoni and M. riccartonensis Zones, yielding abundant
fossils of the Niorhynx Community of Havlíček and Štorch, 1990. In the
Svatý Jan area (Central Segment, Kříž, 1991) comparable bioclastic
limestones formed during the C. murchisoni Zone, are exposed. The
329
deposition of such limestones within the cool-water peri-Gondwana
siliciclastics may reflect the warm interval before the positive δ18O
shift reflecting the onset of the glaciation in the Viki core. The
limestone sedimentation would coincide with the early Sheinwoodian sea-level highstand with highest sea-levels in the C. murchisoni
Zone (Loydell, 1998). In both regions of the Prague Basin, there is only
limited sedimentological information after the C. murchisoni or M.
riccartonensis Zone due to intense volcanism (e.g. Bouček, 1937;
Vaněk, 1962; Kříž, 1991).
For the expression of the Sheinwoodian glaciation, a section at
Lištice near Beroun is of special interest. The Silurian succession here
includes a limestone unit which can be correlated with the Sacta
Member in Bolivia and which was re-deposited during a time of a major
sea-level drop, presumably during the Sheinwoodian glaciation.
In this section, Horný (1955) documented that the irregular
surface on top of the graptolithic shales of the Oktavites spiralis Zone is
overlain by a poorly sorted megabreccia composed of clasts of
organic-rich limestone, and overlain by tuffitic rocks of questionable
age. The clast sizes range from several centimetres to several
decimetres (Horný, 1955: Tabl. 4, Figs. 1, 2). Horný (1955) referred
to two comparable locations in this area and discussed a late C.
murchisoni to Monograptus belerophus Zone age for the deposition of
these limestone beds (equivalent to the uppermost P. amorphognathoides or P. bicornis to lower O. s. rhenana Zone, Figs. 1, 4). We
assume that the gap between the underlying shales and the limestone
breccia is caused by a time of erosion leading to reworking and
transport of the limestone clasts during falling sea-level. Consequently, it is true that this erosional event would pre-date the upper O. s.
rhenana Zone and, like the re-deposition of the Sacta Limestone in
Bolivia, this event correlates to the regression at the beginning of the
Sheinwoodian glaciation (Figs. 3 and 4).
6. Conclusions
(1) New oxygen isotope data from Estonia contribute important
information to clarify the climatically induced extinctions
across the IBE. The δ18O record starting in the Telychian
illustrates significant changes in sea surface temperature in the
Baltoscandian Basin during the studied time interval, but they
show no evidence for a glaciation in the late Telychian Pterospathodus amorphognathoides Zone as proposed by Cramer and
Saltzman (2005).
(2) There is a time offset between the shift in δ13C and the shift in
δ18O suggesting that δ13C data must be used with caution for
interpretations of palaeoclimate.
(3) The interval of the IBE starts in a relatively warm climatic
period prior to the Sheinwoodian glaciation. Cooling starts in
the uppermost part of this biotic event, post-dating several of
the most severe extinctions in different faunal groups.
(4) Our interpretation of global sea-level changes across the
studied interval (cf. Loydell, 2007) contradicts most other
published sea-level curves. In contrast to Cramer and Saltzman
(2005) we suggest a strong regression during the ESCIE.
(5) We suggest that the stratigraphic gap spanning the upper
Ozarkodina sagitta rhenana Zone through a lower part of the
Upper Kockelella walliseri Superzone in the Viki core corresponds a glacio-eustatic sea-level fall as a consequence of
the Sheinwoodian glaciation, as indicated by the highest
δ18O brachiopod values in the Gotland record (Bickert et al.,
1997; Samtleben et al., 2000; Munnecke et al., 2003).
(6) In general, faunas are affected more by global warming and
unstable climate conditions than by the following glacial
events. In addition, cool and stable climates during glacial
periods seem to provide better environmental conditions
reflected by times when faunas flourish and no extinctions
are observed.
330
O. Lehnert et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 296 (2010) 320–331
(7) Since the timing of stable isotope excursions does not coincide
with the IBE, the term ‘Ireviken Event’ should be restricted to
the time slice corresponding to the faunal changes and
stepwise extinctions at distinct datum points as documented
in detail by Jeppsson (1997a,b).
(8) Future studies are needed to clarify whether there is a strong
relationship of climate cooling to certain volcanic events or
event intervals in the early Palaeozoic.
Acknowledgements
The careful reviews by Richard J. Aldridge (Leicester) and one
anonymous reviewer strongly improved the manuscript which is
greatly appreciated. We want to especially thank Mrs. Daniele Lutz
(University of Erlangen) for her help with sample preparation and δ18O
analysis. This project was financially mainly supported by the Deutsche
Forschungsgemeinschaft (Bu-312/44-1). The manuscript was compiled
by OL during his stay at the Czech Geological Survey in Prague (Czech
Republic) in the frame of the “Nachkontakt-Programm” of the
Alexander von Humboldt Foundation (Bonn, Germany), and the
support by the Humboldt Foundation is also greatly acknowledged.
PM thanks the Estonian Science Foundation for the support of his
research (Grant no. 7138), MC acknowledges the Swedish Research
Council for continuous support and JF the Program Kontakt for his
research grant (ME08011). This paper is a contribution to IGCP 503.
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