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Journal of Biogeography
SUPPORTING INFORMATION
The De Geer, Thulean and Beringia routes: key concepts for
understanding early Cenozoic biogeography
Leonidas Brikiatis
Appendix S1: The De Geer, Thulean, and Beringia routes: early and
recent geotectonic evidence
THE DE GEER ROUTE
Etymology
The Swedish geologist Baron Gerard Jacob De Geer (1858–1943) is known for his
papers establishing the glacial isostatic origin (see Mörner, 1979). In 1926, De Geer
first referred to the existence of a physiographic lineament following the western
continental slope of the Barents Shelf from Norway to Svalbard and the north-eastern
continental slope of Greenland and Canada (see Horsfield & Maton, 1970).
Subsequent authors suggested that extensive dextral strike-slip movements occurred
along this lineament during the continental drift of Greenland and North America
away from Eurasia. Harland (1969) named the fault system between Greenland and
Spitsbergen Island (the Spitsbergen fracture zone) the ‘De Geer Zone’. In
biogeography, Szalay & McKenna (1971) first introduced the concept of the De Geer
route. They concluded that:
…if a North Atlantic dry-land dispersal route existed, it lay north of the Norwegian Sea,
crossing the De Geer dextral transform fault zone (Spitsbergen fracture zone) from Greenland
to a juxtaposed Spitsbergen and from there to the rest of Europe via an elevated Barents shelf.
(Szalay & McKenna, 1971, p. 283)
McKenna (1971) coined the term ‘De Geer route’ by referring to the ‘De Geer North
Atlantic dispersal route’ (see also McKenna, 1975 and the references therein).
Early evidence
Simpson (1947) proposed the notion of the dispersal of land faunas across the early
Cenozoic Holarctic landmasses; the exact dispersal mechanisms did not become
known, however, until the discovery in the late 1960s of sea-floor spreading and the
subsequent acceptance of plate tectonics (Vine & Matthews, 1963; Le Pichon et al.,
1973). The initial evidence for a northern Atlantic land corridor was palaeontological
data showing that in the early Eocene, the mid-latitude floras and mammalian faunas
of Europe and North America shared over half of their genera with one another
(Savage, 1971; McKenna, 1975). Beginning in the 1970s, new geological and
geophysical data added tectonic foundations supporting the existence of not only one,
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but two northern Atlantic land bridges: the Thulean route and the De Geer route
(Strauch, 1970; Szalay & McKenna, 1971; McKenna, 1975, 1983a,b; Tucholke &
McCoy, 1986).
According to Szalay & McKenna (1971), the De Geer route was the main
terrestrial passage between Europe and North America; because the Thulean route,
albeit southerly, ‘would have involved island-hopping at best’ and ‘would not suffice,
however, for the large scale faunal transfer that took place in the Sparnacian’ (early
Eocene) (Szalay & McKenna, 1971, p. 284). The contribution of the De Geer route to
biogeographical evolution was even greater than that of Beringia, because ‘the Bering
area would have been about 8 degrees farther north than the Greenland-Barents shelf
route’ (Szalay & McKenna, 1971, p. 284). The De Geer route is suggested to have
operated during the Ypresian (early Eocene) and ended when the land bridge broke
apart near the De Geer fault zone (Spitsbergen fracture zone) at the end of the period
(45–48 Ma) (Szalay & McKenna, 1971; McKenna, 1975).
Szalay & McKenna (1971) and McKenna (1975) considered the Barents Shelf
to be generally subaerial in both the Palaeocene and a large part of the Eocene. The
terrestrial continuity of the Barents Shelf was interrupted in the Spitsbergen fracture
zone, separating North America and Greenland on one side from the Spitsbergen–
subaerial Barents Shelf–Fennoscandia on the other. McKenna (1983a,b) maintained
the same view, considering Fennoscandia as an ‘outpost of North America’ that was
separated from western Europe south of the Baltic by marine barriers until the end of
the Eocene. On the other hand, McKenna also presented updated data supporting a
single subaerial Thulean route [i.e. a more continuous and habitable land connection
from France and the British Isles to North America via the Greenland-Scotland Ridge
(GSR)]. Thus, contrary to Szalay & McKenna (1971), McKenna (1983a,b) considered
the Thulean route to be a more significant biogeographical junction than the De Geer
route, given its more southerly geographical position. McKenna further suggested that
the two land bridges could have simultaneously connected North America to different
parts of Europe (i.e. Great Britain and Scandinavia) near the end of the Palaeocene
and up to some point very early in the Eocene.
Tiffney (1985) generally accepted the palaeogeography proposed by McKenna
(1983a). He kept some reservations, however, concerning whether Fennoscandia and
western Europe remained in isolation in the early Eocene. Furthermore, he noted that
the De Geer route was probably presented as early as the Danian (early Palaeocene)
and maintained until the latest Eocene.
Recent geotectonic and palaeogeographical evidence
Today, the Nares Strait separates Greenland and Arctic Canada (Fig. 2a in the main
paper). It is believed, however, that Greenland was connected with Ellesmere Island
and north-eastern North America in the early Palaeocene (Dawson et al., 1975;
McKenna, 1983a; Blythe & Kleinspehn, 1998; see also Marincovich, 1994;
Golovneva, 2002). Indeed, despite the debate concerning the tectonic structure of the
Wegener Fault, which runs through the Nares Strait (e.g. Tessensohn et al. 2008),
only terrigenous sediments [Eureka Sound Group: Maastrichtian (?) to Eocene strata;
Miall, 1986] have been recorded in the southern part of this area during the latest
Cretaceous–early Palaeocene (Lee et al., 2008, and the reference therein; see also the
land–sea boundary in Ricketts, 1986: Figure 39.1).
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Palaeogeographical reconstructions show another marine corridor to the west,
however, that could constitute a barrier between Greenland-Ellesmere Island and the
north-western Canada landmass (see reconstructions of Torsvik et al., 2002, at 60 Ma;
Stampfli & Borel, 2004, at 70 Ma, Appendix 3 in CD-ROM; Smelror et al., 2009, in
the early Cenozoic). This corridor starts from Baffin Bay, crosses Lancaster Sound,
and reaches as far west as Viscount Melville Sound and the North Canada Basin (see
Fig. 3b). In any case, the stratigraphy west of Lancaster Sound does not support such
a marine continuity. There (as suggested by strata preserved on Devon Island),
Mesozoic sediments of the marine Kanguk Formation (Cenomanian–Campanian),
which is widespread across the Canadian margin, are apparently overlaid by the
terrestrial Expedition Fiord Formation (Maastrichtian–lower Palaeocene), which is
preserved locally in fault-bounded troughs on Devon Island (Witkowski et al., 2011,
and references therein). These strata result in the lower part of the Expedition
Formation of the Eureka Sound Group of Ricketts (1986, 1991, 1994), which is
generally considered to be terrestrial. Miall et al. (1980) correlated these strata with
strata in the Eclipse Trough and the Bylot Island Formation, thus attributing a more
regional range. Smith et al. (1989) noted that such terrestrial sediments, possibly a
lacustrine unit, are found only in the western Lancaster Sound Basin and were either
not deposited in the east or were completely eroded beneath a regional unconformity
above which ‘Tertiary sediments’ were deposited (see stratigraphic charts of Smith et
al., 1989, and Brent et al., 2013). These sediments can be attributed to the Selandian
transgression (at 61.8 Ma; see the sea-level curve in Fig. 4) that probably transgressed
the entire Lancaster Sound Basin interrupting the terrestrial connection (see Fig. 3b).
To the east, Svalbard was connected to Greenland during the latest
Cretaceous–early Palaeocene (Worsley et al., 1986; Muller and Spielhagen, 1990;
Blythe & Kleinspehn, 1998; see supporting floristic evidence in Sweet, 2008, and
references therein, and faunal evidence in Lüthje et al., 2010). Since the late
Palaeocene, a shallow corridor may have separated the two areas during sea-level
highstand (see Fig. 3b). Τhe break up of Svalbard and Greenland started in the latest
Eocene (anomaly 13, c. 36 Ma: Schluter & Hinz, 1978) with a shallow marine setting.
A continuous deep-water oceanic corridor was not established until the opening of the
Fraim Strait in the Neogene (Engen et al., 2008).
Eastwards of Svalbard, the Barents Sea served as part of a marine passageway
connecting the North Sea and northern Atlantic Ocean to the Arctic Ocean during
much of the Mesozoic and Cenozoic. Because of a lack of evidence and the
complicated tectonic history of the region, however, the exact palaeogeographical
regime of the Barents Sea remains unclear for certain periods. Apart from the southwestern Barents Sea, where deeper basins formed by subsidence since the Early
Cretaceous, the rest of the Barents Shelf remained shallow and more tectonically
stable since the Late Carboniferous (c. 300 Ma) (Faleide et al., 1993) (Fig. 2b).
Nevertheless, several parts of the floor of the Barents Sea may have uplifted and
subsided more than once in the Cenozoic. Today, the upper section (most of the
Palaeogene section and the entire Neogene section) of the sea floor is absent or eroded
in most of the Barents Sea drilling wells (Henriksen et al., 2011). On the other hand,
the timing of the major uplift and erosion periods has been a subject of discussion for
many years, because different areas of the Barents Sea achieved maximum burial at
different times (Henriksen et al. 2011). During the Pliocene–Pleistocene, the entire
Barents Shelf was eroded and large amounts of sediment were shed towards the shelf
margin (Smelror et al., 2009). As a result, Cenozoic sediments in the Barents Sea are
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restricted to the south-western basin areas and the western and northern passive
continental margins bounding the shelf.
In the south-western Barents Sea, the base of the Torsk Formation lies
transgressively upon Mesozoic strata, which range in age from Triassic to
Maastrichtian (Ryseth et al., 2003; Nagy et al., 2004). A major stratigraphic hiatus is
present everywhere in the Barents Sea, spanning from the Maastrichtian (latest
Cretaceous) to the late Danian (early Palaeocene) (Faleide et al., 1993; Nagy et al.,
2004; Setoyama et al., 2011) (see Fig. S1.1 below). Danian (early Palaeocene)
deposits are generally rare except in the south-western basin areas (Torsk Formation,
Bjørnøya Basin, Senja Ridge-Vestbakken Volcanic Province, Sørvestsnaget Basin,
Tromsø Basin and Hammerfest Basin) (Ryseth et al., 2003; Nagy et al., 2004;
Setoyama et al., 2011) (Fig. 2b). In contrast, late Palaeocene (Thanetian) deposits
exist in the western, south-western and other parts of the Barents Sea. According to
Ryseth et al. (2003), biostratigraphy and log patterns indicate that these depositions
seem to form a laterally continuous facies, while their occurrence throughout the
Barents Sea supports the idea of Nøttvedt et al. (1988) of a widespread late
Palaeocene epeirogenic sea in the region. Nøttvedt et al. (1988) inferred that a broad
subsidence in the Barents Shelf and the continental shelf of northern Greenland
produced a large epicontinental marine basin during the late Palaeocene. This
epicontinental setting, however, terminated in the early Eocene because of rifting and
volcanism associated with crustal break-up (Nøttvedt et al., 1988; Ryseth et al.,
2003). In such a scenario, a formerly subaerial phase of the Barents Shelf (Fig. 3a)
could have been transgressed and drained through the south-western Barrents Sea
basins (Fig. 3b,c), suggesting that the latest Cretaceous–earliest Cenozoic hiatus in the
continuity of marine depositions probably corresponds to a subaerial exposure of the
shelf.
A regional subaerial exposure of the areas surrounding the Barents Sea is also
predicted by the model of Lyberis & Manby (1993), which constitutes one of the
major current structural hypotheses explaining the formation of the West-Spitsbergen
Fold-and-Thrust Belt. Under this hypothesis, latest Cretaceous–Palaeogene
intercontinental compressional tectonics in the Greenland–Svalbard margin caused
compressive orogen resulting in regional uplift of the Barents Shelf area. On the basis
of foraminiferal assemblages from five wells drilled in the south-western Barents Sea,
a recent study on the palaeobathymetric trends of the area concluded that significant
Late Cretaceous–Palaeocene uplift took place prior to the break-up of the Greenland–
Norwegian Sea. Specifically, Setoyama et al. (2011) estimated uplift and subsidence
rates being between 0.11 mm/year [1000 m/Maastrichtian (70.6 Ma)–Danian (61.1
Ma)] and 0.56mm/year [4000 m/late Maastrichtian (69.1 Ma)–early Danian (62 Ma)]:
results that are compatible with the Lyberis & Manby (1993) model. Under these
results, Setoyama et al. (2011) did not reject the perspective that the late
Maastrichtian–early Danian hiatus in the south-western Barents Sea basinal areas is
due to subaerial exposure, even though, on the basis of morphogroup analysis of the
foraminiferal assemblages across the hiatus, these basins are generally recognized as
being under deep water during this interval (Nagy et al., 2004; Setoyama et al., 2011).
Therefore, it seems likely that the marine setting in the basinal areas of the southwestern Barents Sea was uninterrupted. But given the observed shallowing in the
regional bathymetry, the surrounding shallow Barrents Shelf was probably exposed
subaerially to the east.
The evidence presented thus far favours the possibility that the observed late
Maastrichtian–early Danian stratigraphic hiatus (Fig. S1.1) corresponds to a subaerial
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exposure of the Barents Shelf. Some of the current palaeogeographical reconstructions
of the region seem to be congruent with such perspective: the reconstruction of the
epicontinental seas and straits in northern Eurasia (Akhmetiev & Beniamovski, 2009:
Figure 3), the palaeogeographical reconstruction of the Tethys Ocean c. 70 Ma by
Stampfli & Borel (2004: Appendix 3 in CD-ROM), and the early ‘Tertiary’
(Palaeocene) reconstruction of the northern Atlantic by Ziegler (1988).
The continuous facies deposited locally since the late Danian, but regionally
since the Selandian, throughout the western and south-western Barents Shelf could be
interpreted as the transgressive remains of the exposure of a seaway connecting the
Norwegian-Greenland Seaway to the Arctic Ocean. Indeed, such a seaway is included
in many recent palaeogeographical reconstructions (e.g. since the mid-Danian,
Akhmetiev & Beniamovski, 2009: Figure 3; late Palaeocene, Akhmetiev et al., 2012:
Figure 1b; c. 60 Ma, Torsvik et al., 2002; in the late Thanetian, Brunstad et al., in
press: Figure 4; in the early Cenozoic, Smelror et al., 2009; c. 50 Ma, Brinkhuis et al.,
2006; Gleason et al., 2009; Eberle & Greenwood, 2012). All these reconstructions run
absolutely contrary to McKenna’s (1983a) view that the De Geer route persisted
during the latest Palaeocene and early Eocene.
The available stratigraphic evidence, which allows direct observation of the
past geology, is limited to the southern portions of the Barents Sea floor and extends
as far east as the Nordkapp basin (Nagy et al., 2004) (Fig. 2b). This evidence suggests
that a marine barrier persisted to the north of Fennoscandia from the late Palaeocene
at least until the middle Eocene, but it cannot support the existence of such a barrier
eastwards of the Nordkapp basin (see the palaeogeographical reconstruction of the
Eocene Barents Sea in Smelror et al., 2009, p. 123). This is because very little is
known about the Palaeocene and Eocene palaeogeography of the eastern parts of the
Barents Sea and the Pechora and Kara seas (Fig. 2a). Current geological evidence for
these areas suggests that they probably constituted a tectonically stable epicontinental
mega-region and were either uplifted continental hinterlands or shallow marine seas
with very limited net deposition (Smelror et al., 2009). The sediments that may have
been deposited would have been subsequently removed due to later Neogene uplift
and erosion. Consequently, there is very little direct geological evidence remaining to
help establish models of the Palaeocene and Eocene palaeogeography of the Barents
Sea platform and the Kara Sea (Smelror et al., 2009). Nevertheless, as discussed
below, the remaining stratigraphic sequences include sufficient marine faunal
evidence, albeit fragmentary, to improve our knowledge of the true early Cenozoic
palaeogeography of the Barents Shelf.
Based on the marine faunas, there is no palaeontological evidence for a latest
Cretaceous/early Palaeocene marine connection either between the Arctic Ocean and
the northern Atlantic Ocean or between the Arctic Ocean and western Greenland
(Baffin Bay) (Marincovich, 1994). In contrast, although marine faunas in the Arctic
Ocean during the Danian were geographically confined, there was an excellent
interchange of marine molluscs and microbiota between the type Danian Denmark
fauna and the Danian faunas of western Greenland (Marincovich et al., 1990). The
fact that these more southerly and very diverse Danian faunas shared no mollusc
species with the Arctic Ocean fauna suggests that the Arctic Ocean was isolated at
that time (Marincovich, 1993, 1994). On the other hand, the appearance of the bivalve
Cyrtodaria rutupiensis in the Thanet Sands (type Thanetian) of the London Basin
(Strauch, 1972) confirms a late Palaeocene marine connection between the Arctic
Ocean and the southern North Sea Basin. Cyrtodaria rutupiensis dwelled only in the
Arctic Ocean during the Danian, suggesting that a land barrier existed at that time
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between northern Europe and northern Greenland (Marincovich, 1993, 1994). The
Thanetian first appearance of the bivalve Gari (Garum) in western European faunas is
further evidence of a northern marine connection, because the oldest records of both
Cyrtodaria and Gari appear in the Danian fauna of the Prince Creek Formation in
northern Alaska (Marincovich, 1993).
Nagy et al. (2000) reported the recognition of three agglutinated foraminiferal
species from the late Palaeocene in the Firkanten Formation of Spitsbergen:
Reticulophragmium arcticum, Reticulophragmium boreale, and Labrospira turbida.
All three species are known from the lower part of the late Palaeocene in the
Beaufort-Mackenzie Basin (Arctic North America) (McNeil, 1997) and the late
Palaeocene in Hole M0004A at Lomonsov Ridge (Arctic Basin) (Labrospira sp. in
the latter) (Expedition 302 Scientists, 2006). Furthermore, the species
Psammosphaera eocenica as well the genus Verneuilinoides are shared among the
Spitsbergen and Lomonsov Ridge faunas. The above similarities provide clear
evidence of marine faunal exchanges between the western Barents Sea and the Arctic
Ocean via the Barents Shelf in the late Palaeocene.
The genus Reticulophragmium is believed to have evolved in the Northern
Hemisphere during the ‘mid’-Palaeocene and is not known from Cretaceous strata.
Primitive species of Reticulophragmium are also known from the Palaeocene strata of
the North Sea and western Siberia (Expedition 302 Scientists, 2006). Thus, based on
the occurrence of Reticulophragmium, Verneuilinoides, and other Boreal-Arctic
foraminifera from the western Barents Sea (Torsk Formation), the central Spitsbergen
Basin, the Turgai Strait, the western Siberian Basin, and the Polar Arctic, Akhmetiev
et al. (2012) concluded that a meridional seaway system between the Arctic and
Tethys oceans connected northward to the northern Atlantic Ocean via the Barents
Sea (see the late Palaeocene reconstruction in Akhmetiev et al., 2012: Figure 1b). The
system started in western Asia in the Late Cretaceous; it was interrupted around the
Cretaceous/Palaeogene (K/Pg) boundary and resumed from the mid-Danian to the
Lutetian (Akhmetiev et al., 2012). The interruption corresponds to the period during
which the De Geer route should have been exposed (see Discussion).
THE THULEAN ROUTE
The Thulean land bridge consisted of the subaerial exposure of the shallow transverse
GSR, a bathymetric sill that hampered the deep-water flow between the NorwegianGreenland Sea and the northern Atlantic, extending from south-eastern Greenland,
through Iceland, to the UK continental shelf (e.g. Thiede & Myhre, 1996) (Fig. 2c).
The word Thule comes from ancient Greek literature and refers to an island or region
of the North Sea that has been associated with Iceland (Sprague de Camp, 1954). The
first direct geological evidence of the existence of the Thulean land bridge came from
the recovery of subaerially erupted basalt lavas from the northern flank of the IcelandFaeroe Ridge [Deep Sea Drilling Programme (DSDP) borehole 336; Shipboard
Scientific Party, 1976; Nilsen, 1978]. Palaeontologists have long considered the
Thulean route to be important in the biogeographical history of marine and terrestrial
biota (Strauch, 1970, 1972; Szalay & McKenna, 1971; McKenna, 1975, 1983a,b;
Akhmetiev et al. 1978; Friedrich & Símonarson 1981; Axelrod, 1983).
Morphologically, the GSR can be divided into three parts (Thiede & Myhre,
1996): (1) the Greenland-Iceland Ridge, or Denmark Strait; (2) the Iceland-Faeroe
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Ridge; and (3) the Faeroe-Shetland Channel. At the western end of the Thulean landbridge, possible links between Greenland and North America are provided by the
shallow bathymetric sill at the Davis Strait between southern Baffin Island and southwestern Greenland and, alternatively, by the higher latitude land connection via the
Nares Strait (McKenna, 1983a) (Fig. 3d). A route via the Davis Strait would have
certain biogeographical advantages over the Thulean route; this is an issue currently
under discussion, depending on whether continental or oceanic crust underlies the
bottom of the strait (Srivastava & Arthur, 1989; Chalmers & Pulvertaft, 2001). At the
eastern end of the Thulean route, the Dover strait may have played its own role in
keeping terrestrial continuity. Because of limited data, however, the geotectonic
evolution of this area is not well understood (Van Vliet-Lanoë et al., 2004).
The GSR is believed to be part of a mantle plume system centred under the
Icelandic plateau. High heat flow and low mantle densities under Iceland result in
both the subaerial exposure of the plateau and the relatively shallow depths of other
proximal features (Nunns, 1983; Wright & Miller, 1996). In the early Palaeogene,
magmatic events prior to and during continental separation, and the post-breakup
continuous activity of the Iceland melting anomaly, have resulted in one of the largest
igneous provinces in the world: the North Atlantic Igneous Province (NAIP) (White
& McKenzie, 1989; Jolley & Bell, 2002). Following Morgan (1971), most researchers
explain the early Palaeogene volcanism of the NAIP in terms of lithospheric
impingement of the proto-Iceland mantle plume, although the mantle plume concept
is currently being challenged as an explanation for the NAIP, and alternative models
have been suggested (Meyer et al., 2007, and references therein).
The NAIP was emplaced in two main magmatic phases. The first phase
occurred in the ‘mid’-Palaeocene (c. 62–58 Ma) (e.g. Rousse et al., 2007). The second
phase occurred within Chron 24r (late Thanetian to early Ypresian), spanning the
Palaeocene/Eocene boundary, and was associated with regional uplifts and extensive
volcanic deposits that began at the end of the first phase (Saunders et al., 1997). Thus,
the GSR could have existed at its maximum subaerial continuity only during the
second phase of the NAIP emplacement. After its accretion, most parts of the GSR
transverse are believed to have been continuously subaerial at least until the
Oligocene (Fig. S1.2 below) (see also Denk et al., 2011). Then, the ridge subsided
below sea level through a combination of thermal subsidence and erosion.
The onset of the subsidence of parts of the GSR has been linked to the onset of
the exchange of intermediate and deep-water masses (North Atlantic Deep Water
(NADW)) between the Nordic Seas and the Atlantic Ocean across the GSR (e.g.
Hohbein et al., 2012). Whatever the exact time frame for the onset of the NADW, the
linkage between the Faeroe-Shetland channel and the Faeroe Bank channel
constituted the main passageway (Stoker & Varming, 2011). The other significant
deep-water passageway across the GSR, the Denmark Strait, may also have developed
at about the same time as the GSR became fully submerged (Thiede & Eldholm,
1983).
In contrast to the other components of the GSR, the Faeroe-Shetland Basin
was preserved as a marine property in the early Palaeogene; and as Jones (2011: Box
3.7) notes, according to the evidence available to McKenna (1983a), the land
connection of the Thulean route would have extended only from Greenland to the
Faeroes during the mammalian dispersal interval. The Faeroe-Shetland Basin between
the Faeroes and Scotland would have been a seaway several tens-of-kilometres wide
and several hundred meters deep, constituting a significant barrier to the dispersal of
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land mammals. It was only recently that three-dimensional seismic data combined
with drilling-well data from the southern Faeroe-Shetland Basin provided evidence of
a closure of the former seaway (Smallwood & Gill, 2002; Shaw Champion et al.,
2008; Hartley et al., 2011), and thus, of the establishment of a land connection
between North America and Europe at precisely the time of the observed mammalian
dispersal (c. 56-54 Ma according to the most recent calibrations of the North
American land mammal ages (NALMAs) and European land mammal ages (ELMAs)
against the geomagnetic time-scale of Gradstein et al., 2012). Therefore, the southern
Faeroe-Shetland Basin has been recognized as the gateway for the Thulean route.
In the following section, the correlation of stratigraphy from the southern
Faeroe-Shetland Basin with independent data suggests that the first exposure of the
Thulean route took place in at least two episodes: c. 57 Ma and c. 56 Ma (see Fig.
S1.3 below).
Stratigraphy of the Southern Faeroe-Shetland Basin
The generalized stratigraphy of the early Palaeogene in the wells in Quadrant 204 of
the southern Faeroe-Shetland Basin (Fig. 2d) reveals that the marine property was
interrupted by hiatus or terrestrial intervals (e.g. Shaw Champion et al., 2008). Within
the quadrant, the lower Eocene is best developed in the Judd Basin area, where up to
about 750 m of sediment were deposited (Sørensen, 2003). Wells such as the British
Geological Survey borehole 99/03 (Fig. 2d) are considered to be representative of the
regional stratigraphy (Stoker & Varming, 2011; Stoker et al., 2012).
The unconformities considered below fall into the time interval following the
first magmatic phase of the NAIP, when a possible subaerial exposure of the Thulean
route would be possible (i.e. after c. 60 Ma). Figure S1.2 identifies the exact ages of
these unconformities on the basis of the biostratigraphic markers presented by Shaw
Champion et al. (2008: Figure 2) and the recent bio-chronostratigraphic chart of
Vandenberghe et al. (2012). Thus, the first regional unconformity (known also from
the North Sea; Mudge and Jones, 2004) is dated 59.2 Ma on the basis of the first
appearance of the dinoflagellate cyst Areoligera gippingensis. This hiatus marks the
base of the Lamda Formation, which marks a change from the deposition of deepwater turbidites to a deltaic succession and the first major progradation of the
shoreline into the basin during Thanetian times (Shaw Champion et al., 2008).
Although this unconformity reduced the distance between the Shetland and Faeroe
land margins (coinciding with the PL4 sea-level low stand shown in Fig. S1.3), it
probably did not manage to accrete the Thulean land bridge in full.
A second unconformity appears at the end of the Lamda Formation and the
base of the lower Flett Formation. The latter is correlated with the lower Forties
Formation of the North Sea (Mudge & Bujak, 2001), the base of which is dated to the
upper Thanetian (‘intra-Upper Thanetian’: Mudge & Jones, 2004). Therefore, the
second unconformity falls into this age. Smallwood & Gill (2002) showed that this
unconformity is marked by a surface upon which a branching network of valleys and
intervening topographic highs is preserved. They also identified three valley systems
that drain northward into the Judd Basin, and they interpreted the unconformity to
have formed by subaerial erosion. Shaw Champion et al. (2008) identified the same
unconformity in a wider area, noting that an unconformity of this age had also been
identified across the deeper parts of the Faroe-Shetland basin and in the northern and
central North Sea basins. Therefore, the dating of this unconformity is of special
importance because it probably signals the exact time of the first full exposure of the
9
Thulean land bridge. According to stratigraphic column of Shaw Champion et al.
(2008: Figure 2), this ‘intra-Upper Thanetian’ unconformity falls within the interval
of the first appearances of Apectodinium augustum (56.5 Ma) and Apectodinium
nomomorphum (57.2 Ma). Brunstad et al. (in press) attributed this unconformity to
the Lista-Sele facies transition, which reflects a dramatic sea level fall, as seen
throughout the UK shelf and onshore, and a simultaneous basin restriction with anoxia
in the central and deeper parts of the North Sea basin. This sea-level lowstand interval
coincides with the PL6 sea-level lowstand shown in Fig. S1.3 (a 56.6–57 Ma hiatus;
Kominz et al., 2008: supplementary material). Hence, this interval likely coincided
with the hiatus spanning the ‘Intra-Thanetian unconformity’ and corresponds to the
first full exposure of the Thulean land bridge (Fig. 3d) (see also the
palaeogeographical reconstructions of Stoker & Varming, 2011: Figure 90).
This late Thanetian terrestrial phase was terminated by the transgressive lower
Flett Formation (Figs S1.2 & S1.3), which was largely deposited beyond the Lamda
Formation shelf edge, resulting in shelf bias, erosion, and non-deposition along the
southern margin of the Faeroe-Shetland Basin (Stoker & Varming, 2011). The Flett
Formation is missing in this area, and the younger Balder Formation rests
unconformably on the Lamda Formation (Smallwood & Gill, 2002) because of
erosion. Although marine dinoflagellate cysts, especially the genus Apectodinium, are
abundant in the lower Flett unit (corresponding to the T40 sequence of Lamers &
Carmichael, 1999), terrestrial forms dominate assemblages in the overlying upper
Flett unit (T45 sequence) (Shaw Champion et al., 2008; Stoker & Varming, 2011).
The boundary between the lower and upper units of the Flett Formation is well
restricted by the last occurrence of the dinoflagellate cyst Apectodinium augustum
(55.8 Ma). This age matches perfectly a retreat of the sea (the EL1 sea-level lowstand
shown in Fig. S1.3). Thus, the onset of the terrestrial upper Flett Formation likely
resulted from a second subaerial exposure of the Thulean land bridge coinciding with
the EL1 sea-level lowstand shown in Fig. S1.3.
BERINGIA
Beringia, first proposed by Hultén (1937), is hypothesized to have been a land bridge
between Asia and North America during the Plio-Pleistocene. A pre-Quaternary
connection was acknowledged only after the broad acceptance of plate tectonic theory
(Hopkins, 1967; Cox, 1974). More recently, palaeontological investigations of
correlative fossil-bearing rocks in regions of Alaska (Aniakchak National Monument,
numerous localities along the Colville River, and the Denali National Park; see
Fiorillo, 2008 and references therein) and Utah (Cifelli et al., 1997), combined with
revised tectonic reconstructions of the region (Lawver et al., 2002), led to the
conclusion that Beringia originated approximately 100 Ma. This hypothesis is
congruent with other evidence from marine faunas suggesting that the Bering Strait
was closed as early as the late Albian (Jagt-Yazykova, 2012; Iba et al., 2011).
Some authors consider Beringia to have been terrestrial from the late Albian
until the late Miocene (Marincovich et al., 1990; Marincovich & Gladenkov, 1999;
Gladenkov et al., 2002). Because of the obvious faunal similarities found across the
early Cenozoic of Holarctica, Beringia is acknowledged as a possible dispersal route
for mammals (McKenna, 1983a,b; Beard, 1998; Beard & Dawson, 1999). Marine
evidence, however, leads other authors to leave open the possibility of a marine
connection between the Arctic and Pacific oceans in the latest Cretaceous (Jagt-
10
Yazykova, 2011) (Fig. 7) and early Cenozoic (Gleason et al., 2009). Thus, Beringia is
emergent in some palaeogeographical reconstructions (Eberle & Greenwood, 2012)
and submerged in others (Gleason et al., 2009).
On the basis of climatic (Appendix S2 and Fig. 4), floristic, and vertebrate
evidence (Figs 5a,b & 6), two time windows are likely for biotic exchanges across
Beringia during the Palaeocene: Bering route 1 65.5 Ma (Fig. 7d) and Bering route 2
c. 58 Ma. Possible Eocene exposures are not considered here.
Figure S1.1 Lithostratigraphy of the south-western Barents Sea. Modified from
Setoyama et al. (2011).
11
Figure S1.2 Palaeogeographical continuity of the Thulean land bridge. The sections
correspond to the parts of the Thulean land bridge. Horizontal lines note the ages
proposed by various authors concerning the subsidence of each part. Most parts of the
Thulean land bridge persisted subaerially at least until the Oligocene. Note the
exception of the southern Faeroe-Shetland Basin, which was dominated by tectonic
and sea-level changes resulting in the gateway of the Thulean route. Schema based on
Denk et al. (2011). Biostratigraphy and T-sequences are after Shaw Champion et al.
(2008), and chronostratigraphy is after Vandenberghe et al. (2012).
12
Figure S1.3 Correlation of the eustatic sea-level curve from New Jersey with the
stratigraphies of the southern Faeroe-Shetland Basin, the Kangerlussuaq Basin (East
Greenland), the Turgai Strait, and the mammalian biostratigraphy of North America
and Europe. Figure modified from Brikiatis (in preparation).
13
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18
Appendix S2: Palaeo-climatic conditions around the K/Pg boundary
INTRODUCTION
Climate is a major factor affecting the extension, structure, and composition of
bioprovinces. Hence, past climatic oscillations are of special importance for
understanding and interpreting past biotic changes.
Relatively warm temperatures around the K/Pg boundary are proposed for the
peri-Arctic regions. Specifically, in post-Cenomanian times, floral differentiation
among near Arctic regions suggests a relatively warm Arctic Ocean [supplied by heat
transported northwards along the Western Interior Seaway (WIS)], compared with a
relatively cold northern Pacific Ocean gyre (Spicer & Herman, 2010; Zakharov et al.,
2011). Perhaps because of the WIS, the temperature gradient during the latest
Cretaceous has been interpreted to be less steep (0.4 ± 0.1 °C/°latitude) than it is in
the present day (0.6 °C/°latitude). Temperatures in the latest Cretaceous are thought to
have decreased from about 30 °C near the equator to about −5 °C at the poles; and air
temperatures above 30° palaeolatitude are thought to have been higher than at present
(Amiot et al., 2004). During the time interval when dinosaurs were present on the
northern slope of Alaska, the mean temperature in that region is reported to have
ranged from 2 to 4 °C, for the coldest monthly mean, to 10 to 12 °C, for the warmest
monthly mean (Parrish et al., 1987). In general, although the evidence supports an icefree Arctic summer, the presence of intermittent sea ice in the Arctic winter (Davies et
al., 2009; Spielhagen & Tripati, 2009; Spicer & Herman, 2010 and references therein)
agrees with a wide body of evidence suggesting low winter temperatures in the Late
Cretaceous Arctic region (Amiot et al., 2004; Falcon-Lang et al., 2004; Spicer &
Herman, 2010 and references therein), making hypotheses of mean annual
temperatures higher than 15 °C in the region (Jenkyns et al., 2004; Zakharov et al.,
2011) seem doubtful. In any case, today it is well known that the latest Cretaceous
(Maastrichtian)–earliest Palaeoegene (early Palaeocene) period was characterized by
great climatic variability (Wolfe & Upchurch, 1987; Barrera, 1994; Li & Keller,
1998a, 1999; Francis & Poole, 2002; Nordt et al., 2003; MacLeod & Huber 2005;
Twitchett, 2006; Gallagher et al., 2008; Spicer & Herman, 2010 and references
therein; Zakharov et al., 2011; Hunter et al., 2008; Flögel et al., 2011) that may have
been caused by tectonic reorganization of the oceans (Frank & Arthur, 1999) or/and
the effects of the Deccan volcanism (e.g. Keller et al., 2012). Therefore, general
warming and cooling periods are more important than the prevailing mean
temperatures for our understanding of the palaeoecological and biotic changes that
took place in the peri-Arctic region. For the purposes of the current study, the Kominz
et al. (2008) eustatic sea-level curve from offshore and onshore New Jersey (Fig. 4
and Fig. S2.1 below) is used as a proxy for the palaeoclimatic changes that took place
during the Maastrichtian–Palaeocene period.
Because of limitations of the δ18O record (see below), the eustatic New Jersey
sea-level curve is climatically more informative and useful for explaining the biotic
changes that took place in the Northern Hemisphere (Fig. 4). For example, the
discovery of dinosaur remains in the polar regions was unexpected, because dinosaurs
were previously regarded as animals adapted to warm climates. In order to explain the
discovery of dinosaurs at the higher latitudes, researchers proposed previously
unknown differences between the metabolisms of dinosaurs and those of modern
reptiles (Clemens & Nelms, 1993; Golovneva, 2000) and questioned whether the
19
dinosaurs overwintered at higher latitudes (Clemens & Nelms, 1993) or instead
migrated south with the retreating sunlight (Parrish et al., 1987). The climatic model
based on the Kominz et al. (2008) eustatic sea-level curve suggests that dinosaurs
appeared in the peri-Arctic regions during the latest Cretaceous as a result of the
latitudinal extension of their ranges during a significantly warmer period (Fig. 4).
The sea level record from New Jersey as a climatic proxy
Marine stable-isotope records provide the basis for much of our understanding of past
climates. Oxygen isotope records, in particular, have been used to estimate past water
temperatures, ice sheet volumes, and local salinity variations. Separating the effects of
ice volume from those of temperature is very difficult, because despite the relative
stability of the deep sea, changes in deep ocean temperatures also affect deep-sea
benthic foraminiferal δ18O records (e.g. Wright, 2000). For example, the benthic
foraminiferal δ18O records of the past 50 Myr show a 4‰ increase that must reflect
mostly deep-water cooling; only c. 1.0‰ of the increase can be due to changes in ice
volume (Miller et al., 2005; Miller et al., 2011). Thus, the long-term δ18O record of
the last 100 Myr is thought to reflect c. 12 °C of cooling, complicating the use of δ18O
as a proxy for ice volume beyond the Pliocene–Pleistocene (Miller et al., 2011 and
references therein). Alternatively, Mg/Ca ratios have been used to provide a
palaeothermometer that accounts for the temperature component in deep-sea benthic
foraminiferal δ18O records. Because of the large errors associated with the Mg/Ca
approach, however, the δ18O and Mg/Ca records cannot provide an unequivocal
record of sea levels prior to the Pliocene. Therefore, we must look exclusively to the
sedimentary record of sea-level change (Miller et al., 2011 and references therein).
The sea-level record from onshore and offshore New Jersey (Miller et al.,
2005; Kominz et al., 2008) is considered to be a good proxy of past eustatic sea-level
fluctuations, presupposing the accumulation of at least ephemeral ice sheets in the
Antarctic during the greenhouse world of the Late Cretaceous to middle Eocene
(Miller et al., 2008). This supposition is currently under discussion; but it is supported
by recent climatic models (Flögel et al., 2011).
The main principle of the eustatic sea-level theory is the accumulation of
surface water volume in the form of ice sheets at the high latitudes with a rhythm that
follows long-term climate trends. On the basis of the eustatic sea-level record from
New Jersey, Miller et al. (2008) reconstructed the evolution of the Antarctic ice-sheet
during the Late Cretaceous–Cenozoic; whereas Cramer et al. (2011) reconstructed
trends in ice volume and deep-ocean temperature for the past 108 Myr. Here, the
Kominz et al. (2008) sea-level curve is used as a proxy for the large climatic
fluctuations that took place during the latest Cretaceous (Maastrichtian)–earliest
Palaeogene (early Palaeocene). To indicate the reliability of this climatic model, the
major events presupposed and predicted by the model are correlated with independent
data (see Figs 4 and S2.1).
Correlation of the New Jersey eustatic sea-level curve with independent climatic
proxies
The Cretaceous greenhouse climatic mode terminated with a gradual global cooling
from the late Campanian (c. 73 Ma) until nearly the end the Maastrichtian. Today,
however, it is known that the Late Cretaceous cooling trend was interrupted by two
episodes of greenhouse warming. The earlier episode, the mid-Maastrichtian event
(MME), lasted from about 70 Ma to 68 Ma, while the second episode, the late
20
Maastrichtian event (LME), commenced 450 kyr before the K/T boundary and lasted
only 300 kyr (e.g. Abramovich et al. 2007). The events are represented in the New
Jersey eustatic sea-level curve as the CH3 and CH2 highstands (Fig. S2.1).
The long-term cooling trend that culminated in the early Maastrichtian (CL4
sea-level lowstand in Fig. S2.1), here referred to as the early Maastrichtian cooling
(EMC), is also recorded in marine (Barrera, 1994; MacLeod & Huber, 1996; Barrera
et al., 1997; Li & Keller, 1998a, 1999; Zakharov et al., 1999; Friedrich et al., 2004,
2009; Cramer et al., 2009) and terrestrial oxygen stable isotopes (Nordt et al., 2003).
The formation of ice sheets in high southern latitudes is thought to be a consequence
of the EMC (Miller et al., 1999, 2003). Jagt-Yazykova (2011) noted that during the
EMC, most ammonite forms disappeared from the Arctic, possibly as a result of shortlived, subfreezing conditions that occasionally occurred in the Northern Hemisphere
(Jagt-Yazykova, 2011 and references therein). Furthermore, the retreat of the sea from
the Cretaceous Western Interior Seaway (WIS) (Erickson, 1978; Lillegraven &
Ostresh, 1990; Boyd & Lillegraven, 2011) at the boundary of the Baculites
clinolobatus and Baculites grandis zones (70.5 Ma; Ogg & Hinnov, 2012) match
perfectly with the CL4 sea-level lowstand (70.7 Ma).
The cooling period was succeeded by a hyperthermal event known as the
MME (CH3 sea-level highstand in Fig. S2.1). The MME is supported by both the
terrestrial record (Nordt et al., 2003) and the marine record. In the marine record, the
MME encompasses the global extinctions of the inoceramid and rudistid bivalves and
the latitudinal migrations of some calcareous nannoplankton and planktic
foraminifera. The MME is expressed as an increase in the rate of global climate
cooling, a decrease in the global range of benthic foraminiferal δ13C values from
approximately 3% to less than 1%, and an increase in the rate of change of seawater
87Sr/86Sr isotope ratios (Frank et al., 2005 and references therein). Jagt-Yazykova
(2011) noted that such a temperature maximum in the Arctic Basin at the base of the
late Maastrichtian resulted in one of the fastest ammonite radiations during the Late
Cretaceous, reaching the levels of the ammonite radiations of the late Turonian (JagtYazykova, 2011 and references therein). In the Bering area, such a warm period was
recorded in isotopic palaeotemperatures estimated by δ18O values extracted from
ammonoid shells in southern Alaska sampled from the late part of the early
Maastrichtian (Zakharov et al., 2011: Figure 13).
The succession of EMC to MME can be recognized in the Deep Sea Drilling
Programme (DSDP) sites 525A and 21 in the southern Atlantic (Li & Keller, 1998a,
b) and 463 in the equatorial Pacific (Li & Keller, 1999) where, within the long-term
cooling trend of the late Campanian–Maastrichtian, two prolonged deep-water cooling
events are observed at Chron C31r and C30n. These events are ‘associated with more
gradual cooling in surface waters and separated by temporary warming of 2–3 °C’ (at
Chron C31N) in deep waters (Li & Keller, 1998a, p. 82). These cooling events match
very well with the CL4 and CL2 sea-level lowstands of the Kominz et al. (2008) sealevel curve (Fig. S2.1), while the ‘temporary warming’ correlates with the CH3 sealevel highstand.
The CL2 sea-level lowstand was followed by the CH2 sea-level highstand
(Chron late-C30N) and the CL1 sea-level lowstand. Both events are known from the
marine record (Li & Keller, 1998a, 1999; Olsson et al., 2001).
…near the end of the Maastrichtian (Chron C29R), beginning about 400 kyr before the K/T
boundary and lasting about 100–200 kyr, surface and deep waters warmed rapidly by 3-4˚C
and then sharply cooled by 2–3˚C during the last 100–200 kyr of the Maastrichtian.
21
(Li & Keller, 1998a, p. 82)
In the terrestrial record, the sequence CH2-CL1-CH1(PH1) was printed on the
exported mean annual temperatures from leaf-margin analyses in North Dakota (Wilf
et al., 2003); while the intermediate CH2 (Chron late-C30N) warming event was
independently recognized and referred to as the late Maastrichtian event (LME) by
Nordt et al. (2003). Even the last warming trend of the Cretaceous (leading to the
CH1 sea-level highstand shown in Fig. S2.1) is recorded in both the terrestrial (Wilf et
al., 2003) and the marine oxygen stable-isotope record (Li & Keller, 1998b).
The notable warming led to the PH1 sea-level highstand, referred to here as
the early Danian warming (EDW), and is very close to the recently reported
hyperthermal Dan-C2 event (Quillévéré et al., 2008). The EDW is recorded in the
marine oxygen isotope record in almost all DSDP sites (e.g. Cramer et al., 2009, see
the δ18O curve in Fig. 4; see also Tobin et al., 2012 for a recent Antarctic record). In
contrast, the PL1 sea-level lowstand shown in Fig. 4, here referred to as the midDanian cooling (MDC), is not represented in the δ18O record to the extent expected. A
better image of a Danian cooling just after the warming of the K/Pg boundary, based
on planktonic foraminifera and clay mineralogy, is reported from Kazakhstan (Pardo
et al., 1999). Jagt-Yazykova (2011) noted that evidence from the Russian Far East
suggests a generally abrupt and strong climatic cooling during the Danian with annual
temperatures less than 5°C above freezing (Jagt-Yazykova, 2011 and references
therein).
The PH3 sea-level highstand corresponds perfectly with recent estimations of
the latest Danian event (LDE): a hyperthermal event that has been unequivocally
identified in benthic foraminiferal isotopes from shelf sediments in Egypt
(Bornemann et al., 2009) and in deep-sea material from the Pacific Ocean
(Westerhold et al., 2011). Instead of the early estimations for a ‘Top C27n’ age of this
hyperthermal event, recent studies suggest an early Chron C26r (Bornemann et al.,
2009 and references therein) or 61.75 Ma age for this event (Westerhold et al., 2011),
matching perfectly with the 61.8 Ma age of the PH3 sea-level highstand. Recently,
Schulte et al. (2013) compared the LDE with the Palaeocene–Eocene Thermal
Maximum (PETM; 55.8 Ma) and concluded that both hyperthermal events resulted in
significant eustatic sea-level highstands. The arrows in Fig. 4 show the contrasting
impression of how a significant climatic event is not obvious in the mean values of the
δ18O record but is remarkably clear in the Kominz et al. (2008) sea-level curve.
In general, the Kominz et al. (2008) sea-level model is congruent with the
changes that took place in the palaeofloristic assemblages, the presence-absence of
climatic-indicator taxa such as Cycadophytes (e.g. Golovneva, 2000), and the
anatomy of fossilized woods. Thus, the long-term global cooling that culminated
during the early Maastrichtian (EMC) (CL4 sea-level lowstand in Fig. S2.1) is also
recorded by palaeobotanical (Wolfe & Upchurch, 1987) and fossilized wood-anatomy
studies (Francis & Poole, 2002). During the EMC, the occurrence of the
Cycadophytes-bearing thermophilous vegetation was very rare in the Bering area
(only one flora in the Koryak Upland area); whereas such vegetation was found
everywhere in both northern Alaska and north-eastern Russia during the subsequent
warm period of the MME (CH3 sea-level highstand in Fig. S2.1) (Zakharov et al.,
2011 and references therein). This warming trend is also supported by palaeobotanical
evidence from North America (Wolfe & Upchurch, 1987), the presence of dinosaurs
in the high-latitude peri-Arctic regions, and possibly the occurrence of growth-ring
characteristics in fossilized rare woods in northern Alaska (Spicer & Herman, 2010).
22
Near the Maastrichtian–Danian boundary interval (CL1 sea-level lowstand in Fig.
S2.1), Cycadophytes-bearing, thermophilous vegetation are not found in any of the
Bering floras (Zakharov et al., 2011 and references therein).
Correlation of the New Jersey eustatic sea-level curve with terrestrial climatic proxies
The terrestrial δ18O record does not have the restrictions of the corresponding marine
record in its proxy representation of past climatic fluctuations. It is, however,
governed by another limitation; in the absence of marine faunas, it usually cannot be
well dated. Nordt et al. (2003) reported such a record of δ18O and δ13C values
measured in palaeosol carbonate from North America. A direct correlation with the
Kominz et al. (2008) eustatic sea-level curve, however, encounters two main
problems. First, the record is calibrated on an older time-scale; and second, the two
recognized hiatuses add uncertainty to the time-scale. Figure S2.1 attempts to recalibrate the record of Nordt et al. (2003: Figure 2) with the Gradstein et al. (2012)
geological time-scale and the Kominz et al. (2008) eustatic sea-level curve. For this
propose, the three sections A, B, and C of the Nordt et al. (2003: Figure 2) curve are
re-aligned by photo editing as follows:
Section A: Nordt et al. (2003) correlated the uppermost end of their curve to Chron
C28R. The new applied age for this Chron is 64.7 to 65 Ma. A second calibration
point is the new age of the K/Pg boundary c. 66 Ma (previously recognized
stratigraphically and originally set c. 65 Ma by Nordt et al., 2003). A third calibration
point is the alignment of the LME, also recognized by Nordt et al. (2003) (peak 1 in
Fig. S2.1), to peak CH2 of the Kominz et al. (2008) eustatic sea-level curve.
Section B: The peak of the MME that was recognized by Nordt et al. (2003) (peak 2
in Fig. S2.1) aligned to sea-level highstand CH3 of the Kominz et al. (2008) eustatic
sea-level curve. The sequence boundaries (that are recognized as hiatuses by Nordt et
al., 2003) correlate with the sequence boundaries of the Kominz et al. (2008) eustatic
sea-level curve.
Section C: Nordt et al. (2003) set the lowermost section of the curve on the
Campanian-Maastrichtian boundary, originally c. 71 Ma, now c. 72 Ma according to
the 2012 time-scale.
Although the above approach is simplified, it highlights certain similarities
between the terrestrial climate proxies of Nordt et al. (2003) and the Kominz et al.
(2008) eustatic sea-level curve. This congruency, along with the agreement with the
marine proxies mentioned above, suggests that the eustatic sea-level curve from New
Jersey is a reliable global climatic proxy.
23
Figure S2.1 Correlation of the Kominz et al. (2008) eustatic sea-level curve with
independent proxies from terrestrial and marine records. The sea-level curve is
reconstructed with data from Kominz et al. (2008: supplementary material). The
curves of the δ18O and δ13C values were measured from paleosol carbonate from
North America (Nordt et al., 2003: Figure 2) and recalibrated here (see text). The
benthic δ18O isotope curve is from Li & Keller (1999: Figure 2). The grey shadow
reflects the uncertainty in the dating of the curve as originally designed by Nordt et al.
(2003: Figure 2).
24
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27
Appendix S3: Vertebrate evidence for the existence of Northern
Hemisphere land bridges
DINOSAURS
To investigate the exact dispersal course through Beringia or the De Geer route, it
would be of special importance to determine whether a specific taxon that migrated
from Eurasia to North America was recovered from the western or the eastern bank of
the WIS. This is because the WIS separated western (Laramidia) and eastern
(Appalachia) North America. Taxa recovered from the western bank would suggest
dispersal via Beringia; whereas taxa recovered from the eastern bank would suggest
dispersal via the De Geer route. The retreat of the Cretaceous WIS in the early
Maastrichtian (corresponding to the CL4 sea-level lowstand shown in Fig. 4),
however, resulted in the subaerial reconnection of Laramidia and Appalachia
(Erickson, 1978; Lillegraven & Ostresh, 1990; Boyd & Lillegraven, 2011), thus
‘complicating biogeographical interpretations’ (Sampson et al., 2010: e12292). The
retreat of the Cretaceous WIS is probably why research on dinosaur communities
along the WIS indicates low beta diversity and suggests a single dinosaur community
within the entire WIS region in Maastrichtian North America (Vavrek & Larsson,
2010).
At least one genus of (probably) theropod dinosaurs is thought to have been
shared between Europe and North America during the Maastrichtian. The
‘Euronychodon’ portucalensis from Maastrichtian Portugal and the Paronychodon
from Late Cretaceous North America are proposed to be congeneric (e.g. Antunes &
Mateus, 2003; Pereda-Suberbiola, 2009). The true affinities of both taxa are difficult
to determine, however, because the taxa are known only from fossils of teeth
(Antunes & Mateus, 2003; Larson & Currie, 2013). Therefore, although they may be
valid taxa, they cannot yet be taken as reliable evidence of intercontinental dispersal.
The hadrosaurs were a very diverse and successful dinosaur group that
dominated the Laurasian landmasses during the later stages of the Late Cretaceous
(Horner et al., 2004). Latest Cretaceous representatives of the Lambeosaurinae
subfamily have been recovered from all three Laurasian continents: Asia, North
America, and Europe. The closely related taxa suggest land connections in the
Northern Hemisphere (Godefroit et al., 2003; Cruzado-Caballero et al., 2011). Two
representative pairs of these taxa are the Corythosaurus–Olorotitan from Campanian
North America and late (or ‘mid-‘) Maastrichtian Asia (Godefroit et al., 2003; 2012),
and the Parasaurolophus–Blasisaurus from Campanian North America and late
Masstrichtian Europe (Cruzado-Caballero et al., 2010, 2011) (Fig. 6). Prieto-Marquez
& Wagner (2009) demonstrated evidence suggesting that Koutalisaurus kohlerorum
from early Maastrichtian to early–late Maastrichtian (Pereda-Suberbiola et al., 2009)
Lleida Province in north-eastern Spain is most probably the junior synonym of
Pararhabdodon isonensis from the same region (see the review of European taxa in
Pereda-Suberbiola et al., 2009), providing conclusive evidence of the presence of the
Lambeosaurinae in Europe. The same authors proposed that Tsintaosaurus
spinorhinus from Asia forms a clade with Pararhabdodon isonensis, suggesting that
this clade originated in Asia during the middle or late Campanian and migrated to the
Iberian island of the European archipelago during the Maastrichtian.
28
Another related pair of hadrosaurs is Arenysaurus–Charanosaurus from the
Late Cretaceous of Portugal and the Russian Far East (Amur Region) (Godefroit et
al., 2003; Cruzado-Caballero et al., 2011). The phylogenetic analysis of Godefroit et
al. (2012) suggested a closer relationship between Parasaurolophus and
Charanosaurus; their analysis does not, however, include the closely related
European genus Blasisaurus. In any case, even with the different topologies, the
closed relationships among the lambeosaurins of the three continents of the ‘middle’
or late Maastrichtian Northern Hemisphere constitute clear evidence of faunal
exchanges.
Within the sauroloph clade, the genus Saurolophus is known both in western
Canada (Saurolophus osborni, Horseshoe Formation, lower Maastrichtian) and in
Mongolia (Saurolophus angustirostris, Nemegt Formation, lower Maastrichtian)
(Bell, 2011 and references therein). Its sister taxon, Prosaurolophus, is only known in
western North America and is older (late Campanian, 75 Ma; Gates & Farke, 2009
and reference therein), suggesting dispersal from North America to Asia (Godefroit et
al., 2011). Recently, the genus Saurolophus was identified in the Almond Formation
of Wyoming dated c. 72 Ma (Gates & Farke, 2009). This discovery further supports
the anagenesis of the Prosaurolophus–Saurolophus clade and its later dispersal to
Asia via the De Geer route.
The Alvarezsaurids are a group of small maniraptoran theropod dinosaurs.
Derived members of the Alvarezsaurids (the Parvicursorinae) display many derived
features that also occur in some birds, and they were originally interpreted as very
basal avians (Xu et al, 2010 and references therein). Alvarezsaurids have been found
in Asia, South America, and North America. Previous biogeographical interpretations
proposed dispersal from South America to North America and then to Asia (Longrich
& Currie, 2009 and references therein). More recent discoveries show, however, that
Laurasian taxa can be grouped into a separate derived clade, the Parvicursorinae, with
the exclusion of the South American taxa (Xu et al, 2010; 2011). In this scenario, the
‘middle’ and late Maastrichtian representatives in North America (Albertonykus
borealis and some other unnamed specimens; see Longrich & Currie, 2009: Figs 1 &
2) are considered to have dispersed from Late Cretaceous Asian populations (Xu et al,
2011).
On the basis of recovered teeth, Godefroit et al. (2009) recognized the
presence of dinosaur genera such as Dromaeosaurus and Saurornitholestes at
Kakanaut in north-eastern Russia. These genera are also known from the Russian Far
East (Van Itterbeeck et al., 2005) as well as from the Alaskan (Fiorillo & Gangloff,
2000; Fiorillo, 2008) and mid-latitude sites of North America (e.g. Russell & Manabe,
2002).
The North American theropod Troodon formosus is considered to have been
present on both sides of the Bering area (in northern Alaska and in Kakanaut, in the
Russian Far East), supporting dispersal to Asia via Beringia (Zakharov et al., 2011).
This taxon is considered to have been a year-round resident of the Arctic regions that
was adapted to the cooler climate and winter darkness (Fiorillo & Gangloff, 2000).
Nevertheless, Troodon formosus is known from both sides of the WIS in North
America from northern Alaska to Wyoming. Ιts presence in Kakanaut in the Russian
Far East is based on scarce tooth remains and is tentatively referred to as Troodon cf.
formosus. Recently, however, such tooth morphology was concluded to be widely
distributed among the troodontids and hence cannot be regarded as diagnostic at the
generic level (Godefroit et al., 2009 and references therein). In any case, teeth from
29
troodontids, small theropods, and hadrosaurs have been recovered from all known
dinosaur deposits in the peri-Arctic region, i.e. the northern slope of Alaska, Siberia,
north-eastern Russia, and Bylot Island (northern Canada) (Rich et al., 2002; Godefroit
et al., 2009). Contrary to the arguments of Fiorillo (2008) and Zakharov et al., (2011),
the latest Cretaceous dinosaur affinities among the Laurasian continents suggest
dispersal via the De Geer route, because the floristic record shows that Beringia was
probably transgressed and not exposed during the Maastrichtian (Fig. 7c). This
conclusion is congruent with the interpretation of Godefroit et al.
…the development of very different kinds of dinosaur communities during the Maastrichtian
may reflect some kind of geographical isolation between eastern Asia and western North America
during this time, or important differences in climatic or palaeoecological conditions.
(Godefroit et al., 2011, p. 184)
REPTILES
The crocodylians include the clades Crocodyloidea, Gavialoidea, and Alligatoroidea,
which together incorporate all of the current species of crocodiles, alligators, caimans,
and gharials. The crocodyloids are considered to have a Laurasian origin, and their
earliest members are the closely related Prodiplocynodon langi, from a Maastrichtian
terrestrial horizon in the Lance Formation of Wyoming (Mook, 1941), and
Arenysuchus gascabadiolorum, from late Maastrichtian (within Chron C30n) Spain
(Puertolas et al., 2011) (Fig. 6). The gavialoids are also believed to have originated
from Laurasia; the oldest known member, Eothoracosaurus mississippiensis, was
found in the late Campanian or early Maastrichtian Ripley Formation in Mississippi
(Brochu, 2004). By the Late Cretaceous, the species Thoracosaurus neocesariensis
and Thoracosaurus macrorhynchus appeared, respectively, in the latest Maastrichtianearliest Palaeocene of New Jersey and the early Palaeocene of France, Poland, and
Sweden (Zarski et al., 1998; Brochu, 2004 and references therein). Phylogenetic
analyses (Delfino et al., 2005; Puertolas et al, 2011) support the close relationship
between the crocodyloid and gavialoid taxa from North America and Europe and thus
provide evidence for faunal exchange via the De Geer route around the K/Pg
boundary.
Late Palaeocene Europe and North America shared another gavialoid genus
represented, in each continent respectively, by Eosuchus lerichei and Eosuchus minor
(Delfino et al., 2005 and references therein; Brochu, 2006). In Belgium, E. lerichei
was found in the same layer as the gavial-like neochoristodere Champsosaurus, which
is known also from the late Palaeocene of France and the Late Cretaceous–Palaeocene
localites of North America (Matsumoto & Evans, 2010 and references therein). The
proposed mid-Thanetian age of the fossil beds in Belgium (Delfino et al., 2005)
suggests that both Eosuchus and the Champsosaurus probably migrated via the first
exposure of the Thulean route c. 57 Ma (Fig. 3d), rather than via the De Geer route.
The turtle genus Compsemys (otherwise restricted to the Late Cretaceous and
Palaeocene of North America) is known also from the Cernaysian (late Palaeocene) of
France, suggesting dispersal via the Thulean route c. 57 Ma (Godinot & Lapparent de
Broin, 2003; Lyson & Joyce, 2011).
Godinot & Lapparent de Broin (2003) postulated that early Cenozoic
herpetofaunal relations support direct dispersal from Asia to Europe via the Turgai
Strait. They note that contrary to mammals, there is no turtle genus among the early
30
Eocene immigrants that is common to Europe and North America. Therefore, the
turtles that arrived in Europe in the early Neustrian had to come via a route other than
North America, implying possible dispersal from Asia. Such hypotheses, however,
cannot currently be verified at the genus and species level, because most of the
specimens are not complete enough.
The lizard family Agamidae was thought to have reached Europe directly from
Asia. The Agamidae are known from the Late Cretaceous of Asia (e.g. Dashzeveg et
al., 1995) and have been identified in the late Palaeocene of Germany (Weigelt,
1940), where Agamidae indet. was found together with MP6 faunas such as
Arctocyon sp. and Plesiadapis tricuspidens, thus confirming the Cernaysian age of the
fossils. Agamid lizards made their first undeniable appearance in Europe in the early
Eocene (MP7, locality of Dormaal) in the form of a single genus and species:
Tinosaurus europeocaenus (Augé & Smith, 1997). The agamids managed to enter
North America only later, during the Eocene, as demonstrated by the presence of the
genus Tinosaurus in the middle Eocene (Bridgerian) of Wyoming (Gunnell & Bartels,
2001). Thus, the biogeography of the agamids provides strong evidence for terrestrial
access via the Turgai Strait, at least in the earliest Eocene.
Some early papers proposed that the snake family Boidae was present around
K-Pg times in both Europe and South America (Rage, 1984; Le Loeuff, 1991),
suggesting dispersals via the De Geer route. This proposal remains tentative, however,
because the evidence is scarce.
AMPHIBIANS
The neochoristodere Champsosaurus is known from one of the few fossiliforous sites
in north-eastern North America: the mid-Palaeocene (Tiffanian) Roche Percée in
Saskatchewan, Canada, where the oldest representative of the living salamander genus
Cryptobranchus and the family Cryptobranchidae was revealed (C.
saskatchewanensis; Naylor, 1981). Recently, Skutschas (2009) concluded that two
salamander genera coexisted in the Turonian Byssekty Formation in Uzbekistan: the
cryptobranchid Eoscapherpeton and the cryptobranchoid Nesovtriton. According to
this interpretation, although stem cryptobranchoids could also have been present in
Cretaceous North America, the family Cryptobranchidae originated in Late
Cretaceous Asia and dispersed to North America later, around the K/Pg boundary, via
the De Geer route (Fig. 6).
31
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