The relationships between Late Ordovician sea-level changes and faunal turnover in western Baltica: Geochemical evidence of oxic and dysoxic bottom-water conditions

Palaeogeography, Palaeoclimatology, Palaeoecology 271 (2009) 268–278

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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

The relationships between Late Ordovician sea-level changes and faunal turnover in western Baltica: Geochemical evidence of oxic and dysoxic bottom-water conditions Jesper Hansen a,⁎, Jesper Kresten Nielsen b, Nils-Martin Hanken c a b c

Department of Natural Science, Tromsø University Museum, NO-9037 Tromsø, Norway Basin Modelling Department, SINTEF Petroleum Research, S. P. Andersens vei 15B, NO-7031 Trondheim, Norway Department of Geology, University of Tromsø, Dramsveien 201, NO-9037 Tromsø, Norway

a r t i c l e

i n f o

Article history: Received 13 May 2008 Received in revised form 13 October 2008 Accepted 31 October 2008 Keywords: Ordovician Norway Brachiopods Trace elements Sea-level changes Anoxicity

a b s t r a c t The late Sandbian to the early Katian (Late Ordovician) in southeastern Norway is dominated by marine mudstones that contain an exceptional spectrum of macrofauna and trace element geochemistry, recording both abrupt and gradual faunal changes during major sea-level and environmental shifts. This study investigates the relationships between sea-level changes and their influence on the source of clastic material, oxygen levels in the bottom water and faunal changes, with special emphasis on the brachiopod fauna. Trace element ratios indicate that nearly stable upper dysoxic bottom-water conditions prevailed in the northwestern part of the Baltoscandian Sea. The exception is during the major shallowing in the earliest Katian, when there was an abrupt shift to oxic conditions as the sea bottom came within a normal storm wave base. Nonetheless, the pre-shallowing epibenthic fauna (though not the shelly endobenthic) in this area was rich and diverse and two major immigration phases of new brachiopod taxa are seen well before the shallowing. This indicates that the immigration of new taxa was not the result of an increase in oxygen content in the bottom water. More brachiopod genera stayed during the abrupt shallowing and increased oxic level than did during the following gradual return to deeper, dysoxic environments. The major brachiopod immigration phases took place markedly earlier in this northwestern part of the Baltoscandian Sea than in the central part (Eastern Baltic) and possibly also the comparable faunal turnover in Laurentia. The following disappearance of taxa during an early Katian transgressive event coincided with the faunal turnover in the shallow-water environments of the East Baltic. The depositional history was consistent within the investigated mudstone-dominated offshore facies of the Oslo Region, which are comparable to the Central Baltoscandian Confacies Belt of Baltoscandia. The composition of the siliciclastic material which was deposited in the basin was nearly constant through the sequence, though the land area expanded to include ophiolites, expressed by Cr enrichment, coinciding with the major sea-level drop. The Cr enrichment indicates that, by that time at least, the ophiolite complexes north of the Oslo Region became subaerially exposed. This enrichment thus forms a potential geochemical marker horizon representing the basal Katian Stage. © 2008 Elsevier B.V. All rights reserved.

1. Introduction The Baltoscandian fauna underwent a major turnover during the late Ordovician. This was in response to regional climatic and ecological changes and interchange with non-Baltic faunas as the earlier isolated Baltica slowly converged with Avalonia and Laurentia during its northward drift (Jaanusson and Bergström, 1980; Harper, 1986; Harper and Hints, 2001; Ainsaar et al., 2004; Sturesson et al., 2005; Hansen and Harper, 2008). In the East Baltic part of the Baltoscandian Sea, the most significant changes in facies distribution ⁎ Corresponding author. Department of Natural Science, Tromsø University Museum, NO-9037 Tromsø, Norway. Tel.: +47 77 64 50 26; fax: +47 77 64 55 20. E-mail address: [email protected] (J. Hansen). 0031-0182/$ – see front matter © 2008 Elsevier B.V. All rights reserved. doi:10.1016/j.palaeo.2008.10.022

and faunal composition took place in the earliest Katian (Rõõmusoks, 1972; Hints et al., 1989; Kaljo et al., 1996; Nestor and Einasto, 1997; Meidla et al., 1999; Ainsaar and Meidla, 2001; Kaljo et al., 2004). These significant changes have been less well documented in the Oslo Region, Norway, constituting the northwestern part of the Baltoscandian Sea. The main goal of the present study has been to unravel the relationships between sea-level changes, variations in the oxygen content of the bottom water and faunal changes (with the main focus on the brachiopods). In addition, variations in source area of the siliciclastic material have been investigated. For these purposes, eight sections in the Oslo Region were included so as to elucidate both vertical and lateral changes (Fig. 1A). The Oslo Region is well suited to such investigations because the stratigraphy of the brachiopod taxa is

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269

Fig. 1. (A) Location of the investigated sections (world map modified after Riccardi et al. (2007); right map modified after Hansen (2008)). (B) Palaeogeographic map showing the palaeocontours of the Baltoscandian Sea as indicated by the facies belts (world map modified after Cocks and Torsvik (2004); right map compiled after available literature (e.g. Männil, 1966; Størmer, 1967; Cocks and Torsvik, 2005; Hansen and Harper, 2008)). Facies belts: CB — Central Baltoscandian; E — Estonian; L — Lithuanian; LT — Livonian tongue; O — Oslo; S — Scanian. Other abbreviations: F — Fennoscandia Land; M — Moscow Basin; Sa — Sarmatia Land; S–T — Sorgenfrei-Tornquist fault; T — Telemark Land; T–T — Tornquist– Teisseyre fault.

well known and because the sequence including the faunal turnover appears to be more complete than in the East Baltic region. 2. Geological setting and stratigraphy 2.1. Regional geology Palaeomagnetic studies have shown that, during the Ordovician, the palaeoplate Baltica drifted from high southern latitudes towards the equator (Torsvik et al., 1991). In the Middle to Late Ordovician, the engulfing seas, the Tornquist Sea to the southwest and the Iapetus Ocean to the west, slowly narrowed as the palaeoplates Avalonia and Laurentia converged on Baltica (Bruton et al., 1985; Sturesson et al., 2005; Torsvik and Cocks, 2005). The Tornquist Sea crust was subducted beneath Avalonia during its northward drift and explosive vents associated with Andean-type magmatism supplied Baltica with gigantic Sandbian ash falls controlled by westerly Ordovician winds (Pharaoh et al., 1993). During the closure of the Tornquist Sea and the Iapetus Ocean, island arcs and arc basins were developed to the west of Baltica, together with ophiolites and arc-related plutons. These complexes, which are present along nearly the whole strike length of the Scandinavian Caledonides, have ages varying from late Cambrian to late Ordovician (Dunning and Pedersen, 1988). During the late Sandbian and early Katian, the shallow areas in the Baltoscandian Sea located on the western side of Baltica at about 30° S experienced a shift from temperate-water to warm-water carbonate sedimentation, though globally the seas were slowly cooling (Nestor and Einasto, 1997; Cocks and Torsvik, 2004; Trotter et al., 2008). The Baltoscandian Sea was a pericratonic and epicontinental sea with a ramp facies belt pattern in the outer western part and shelf facies in

the east (Kaljo et al., 2004). The deposits have been divided into facies belts characterized by differences in fauna and lithology. They have been termed “confacies belts” and represent, in general terms, an offshore to nearshore transect from Scandinavia to the east Baltic (Männil, 1966; Jaanusson, 1976). The Oslo Region (Fig. 1B) belongs to the NE–SW oriented Oslo confacies belts. It is characterized by more pronounced basement topography than in other parts of the Baltoscandian Basin, leading to marked lateral and vertical lithological changes (Jaanusson, 1976; Bruton and Harper, 1988). The Oslo confacies belts include up to four belts, according to Bockelie (1978). The studied sections lie within the two eastern offshore Asker and Oslo belts, which in the Sandbian and lowermost Katian Stage have a brachiopod fauna comparable with the Central Baltoscandian Confacies Belt (Hansen and Harper, 2008). The Baltoscandian Sea was tectonically stable and subject to very low sedimentation rates except along the south- and northwestern cratonic margins, which became influenced by the closure of first the Tornquist Sea and then the Iapetus Ocean as the palaeoplates Avalonia and Laurentia converged on Baltica (Bruton et al., 1985; Sturesson et al., 2005). The Oslo Region was partly sheltered from the Iapetus Ocean during most of the Ordovician by western land areas or shallows, the “Telemark Land” (Skjeseth, 1952; Størmer, 1967; Bjørlykke, 1974a; Brenchley et al., 1979; Bruton and Owen, 1982; Bruton and Harper, 1988). 2.2. Local geology In this work, we have concentrated our investigations upon the Upper Ordovician Arnestad Formation (upper Sandbian and lowermost Katian) and the overlying Frognerkilen Formation (Katian) defined by

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correlations of the base of the Arnestad Formation are weak due to a high degree of endemism and few biostratigraphically valuable taxa, but seem to correlate with the basal part of the Haljala Stage and not the top of the preceding Kukruse Stage (e.g. Henningsmoen, 1953; Spjeldnæs, 1957; Qvale, 1980; Hansen and Harper, 2006). In Estonia, in the East Baltic, the lower boundary of the Oandu Stage, succeeding the Keila Stage, is defined at the top of a hiatus. There have been slight disagreements regarding whether to place the missing peak lowstand deposits within the uppermost Keila or the Oandu Stage (cf. Kaljo et al., 2004; Nielsen and Meidla, 2004; Kaljo, pers. comm., 2008). Presently it appears that most investigators are in favour of placing the stage boundary at the beginning of the transgressive deposits (Ainsaar et al., 2004; Nielsen and Meidla, 2004; Nielsen, pers. comm., 2008). Thus, according to the correlation scheme made by Nielsen (2004) for the Oslo Region, the lower boundary of the Oandu Stage correlates to the lower half of the Frognerkilen Formation. When comparing with the Kaljo et al. (2004) sea-level curve, we have corrected for the differences in the defined lower Oandu Stage boundary. The boundary between the Sandbian and Katian global stages lies close to the Grimstorp K-bentonites (Meidla et al., 1999; Ainsaar et al., 2004; Kaljo et al., 2004; Leslie et al., 2004; Nielsen, 2004; Nõlvak et al., 2006).

Owen et al. (1990). Both formations are characterized by dark grey shaly deposits and carbonate mudstone layers with only minor lithological variations (Fig. 3). The Arnestad Formation is characterized by shaly deposits interbedded with carbonate mudstone layers and stratabound carbonate nodules. The shaly horizons are commonly 30–40 cm thick while the nodular and bedded limestone horizons are almost invariably less than 10 cm thick. Some of the thin beds of calcareous mudstones change laterally into nodular beds. Close to the top of the formation, the calcareous mudstone beds become thicker and change from a dominance of nodular beds to massive beds. K-bentonites, varying from 1 mm to more than 1 m in thickness, occur throughout the Arnestad Formation (Hagemann and Spjeldnæs, 1955; Bergström et al., 1995). Some horizons in the Arnestad Formation contain scattered small phosphorite nodules and phosphatized fossils such as gastropods and crinoids indicating, at least periodically, low sedimentation rates (cf. Kennedy and Garrison, 1975; Bentor, 1980). An apparent lack of phosphate in the uppermost part of the Arnestad Formation and the overlying Frognerkilen Formation could indicate an upward increase in sedimentation rate (Möller and Kvingan, 1988). The thickness of the formation is somewhat uncertain owing to a combination of tectonism and incomplete exposures. Originally, it was estimated as about 26– 45 m by Brøgger (1887) and Hageman and Spjeldnæs (1955). However, according to our investigations, these earlier investigations underestimated the thickness due to faults and lack of detailed biostratigraphical data. We obtain a minimum thickness of about 90 m. The overlying Frognerkilen Formation is dominated by dark grey nodular calcareous mudstones with interbedded siliciclastic mudstone. There is a gradual increase in the thickness of the intercalating siliciclastic mudstone beds from the middle of the Frognerkilen Formation and upwards. The Frognerkilen Formation thins eastwards from about 18 m in the Asker area to about 7 m at Keyserløkka (Qvale, 1980). Minute pyrite aggregates are common throughout the whole sequence, both in the matrix and in skeletal material. The deposits are generally lacking any sedimentary structures except for four shell beds with shell pavements in the lower part of the Arnestad Formation and small-scale planar and cross lamination in silt beds in the top of the Arnestad Formation. Barring the four shell beds in the lower part of the Arnestad Formation, the fossils show random orientation along the bedding plane, though they are preferentially closer to horizontal than vertical. The shells may occur scattered or in small pockets. In the four shell beds, the shells normally have the convex side up. The shells are often broken, but mechanical rounding of the fractured surfaces is very limited. Bioerosion is rare and has only been observed in a few cases.

The distribution of body and trace fossils was obtained by a combination of field observations (Fig. 1) and examination of 142 bulk samples, 93 fossil samples and 69 polished thin sections. In general, the stratigraphical distance between the bulk samples in each section was 0.5–2.0 m and the size of a sample was 1–2 kg. The rock fragments in the bulk samples were examined under a binocular microscope to count all visible fossils. More than 25,000 brachiopods, among 75,000 fossils, were investigated. Taxonomic description of the brachiopod fauna has been given by Hansen (2008). The samples are presently stored at the Tromsø University Museum, Norway, but are later to be transferred to the Natural History Museum of Oslo, Department of Geology, Norway. Brachiopod assemblages were defined qualitatively through studies of range charts illustrating abundances of each taxon, as well as quantitatively through cluster analysis and calculations of the species density and species diversity indices. In the quantitative studies, the sequence was divided into seven parts defined by where the fauna changed. A complete treatment of the subject is provided by Hansen and Harper (2008).

2.3. Stratigraphy

3.2. Geochemistry

The composite log in Fig. 3 is based on eight profiles within the Oslo–Asker district of the Oslo Region (Fig. 1). The exposures, though small, are good with little tectonic deformation. Grid references and descriptions of these profiles have been given by Hansen (2008). The K-bentonite beds in the Arnestad Formation are excellent marker beds and have been used for correlation within the Baltoscandian area (Hagemann and Spjeldnæs, 1955; Bergström et al., 1995). The lower part of the Arnestad Formation includes several thin bentonite beds of the Grefsen K-bentonite complex, which correlates to the lower half of the East Baltic Haljala Regional Stage. These are succeeded by a pair of cm thick beds named the Sinsen K-bentonites, which correlates to the upper part of the stage. However, the Sinsen K-bentonites are not present in any of the investigated sections, but from other sections they are known to occur 2–2.5 m below the Kinnekulle K-bentonite, which marks the lower boundary of the Keila Regional Stage (Bergström et al., 1995). Above the Kinnekulle K-bentonite follows a pair of cm thick beds named the Grimstorp K-bentonites, which correlates to the lower part of the East Baltic Keila Stage (Bergström et al., 1995). Biostratigraphical

The geochemical data of the studied sections are presented in Table 1. About 10–20 g of sediment from each of 86 mudstone samples were ground in a Siebtechnik mill. Major and trace elements were determined by XRF and ICP-MS analyses of these samples. Some of the samples were analysed by X-ray fluorescence (XRF), Norwegian Geological Survey, under the following conditions: 5.4 g sample and 1.2 g Hoechst wax, Philips PW1480 X-ray spectrometer, Rh X-ray tube. The other samples (0.25 g) were heated in a strong acid digestion (HNO3–HClO4–HF) to fuming and taken to dryness. The residue was dissolved in HCl. The solutions were analysed by ICP-MS, Acme Analytical Laboratories Ltd., Canada.

3. Materials and methods 3.1. Fossils

4. Results 4.1. Trace fossils The deposits are thoroughly bioturbated, resulting in homogenization of the matrix and poor preservation of primary sedimentary

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Table 1 Geochemistry of the shaly mudstones Sample number

Bed number

Section

Formation

Al (%)

Cr (ppm)

Ni (ppm)

Ti (%)

V(ppm)

Zr (ppm)

258 264 270 268 133 257 131 142 254 134 265 144 140 259 266 263 267 260 262 261 282 279 280 281 273 216 114 110 108 109 99 98 111 112 137 79 87 120 119 201 117 204 115 150 200 54 179 72 162 35 93 183 214 33 92 147 80 81 211 236 82 84 250 170 4 86 178 232 91 90 230 71 219 220 229

1 2 3 4 5 5 6 7 7 8 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 45 46 46 47 47 48 48 49 50 51 51 51 52 53 54 55 56 56 56 56 57 58 59 60 61 62 63

Persteilene Persteilene Vollen 1 Persteilene Vollen 1 Persteilene Vollen 1 Vollen 1 Persteilene Vollen 1 Persteilene Vollen 1 Vollen 1 Persteilene Persteilene Persteilene Persteilene Persteilene Persteilene Persteilene Keyserløkka Keyserløkka Keyserløkka Keyserløkka Vollen 2 Vollen 2 Vollen 2 Vollen 3 Vollen 3 Vollen 3 Vollen 3 Vollen 3 Vollen 3 Vollen 3 Vollen 3 Vollen 3 Vollen 3 Vollen 3 Vollen 3 Vollen 3 Vollen 3 Vollen 3 Vollen 3 Vollen 3 Vollen 3 Nakkholmen Fornebu Nakkholmen Vollen 3 Nakkholmen Vollen 3 Fornebu Vollen 3 Nakkholmen Vollen 3 Vollen 3 Vollen 3 Vollen 3 Nakkholmen Ildjernet Vollen 3 Vollen 3 Ildjernet Fornebu Nakkholmen Vollen 3 Fornebu Ildjernet Vollen 3 Vollen 3 Ildjernet Nakkholmen Ildjernet Ildjernet Ildjernet

Arnestad Arnestad Vollen Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Arnestad Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen

8.04 7.53 7.66 7.93 8.14 7.59

334.5 247.3 190.9 240.0 299.7 290.1 336.0 308.9 299.1 318.4 318.4 289.7 280.8 302.3 245.5 259.6 326.2 360.3 304.5 340.0 298.9 375.9 545.7 280.7 337.0 249.2 116.5 361.0 253.7 206.4 205.6 208.5 289.2 510.0 284.0 351.0 466.0 476.0 310.6 242.0 317.0 259.6 206.7 330.1 288.0 371.7 180.0 341.9 342.5 306.1 252.6 278.0 426.0 345.6 271.6 354.4 398.0 623.0 612.5 525.2 675.1 210.0 524.2 265.7 589.7 489.1 830.0 719.0 241.9 260.3 331.6 551.3 547.0 536.0 382.0

200.5 161.5 168.0 195.7 197.5 186.2 206.0 238.9 193.7 244.9 252.4 218.1 203.6 226.5 188.4 204.0 220.3 243.2 212.7 207.0 220.7 243.5 246.8 232.9 161.0 110.7 81.2 166.0 177.5 148.5 145.4 148.8 202.9 218.0 163.0 206.3 211.0 223.0 202.4 129.0 213.5 174.1 181.1 195.7 165.0 207.7 89.0 208.3 211.6 196.5 182.5 166.0 196.0 185.2 170.0 212.7 163.0 219.4 247.9 218.9 260.9 126.0 231.7 186.4 271.5 280.6 311.0 269.0 184.3 142.7 238.2 286.7 267.0 281.0 205.0

0.496 0.487 0.407 0.459 0.440 0.474

187 183 136 167 149 182 198 162 176 165 192 166 148 185 164 167 185 174 156 166 189 203 151 188 139 120 81 198 159 123 121 126 150 186 141 175 209 196 157 100 165 146 158 158 116 191 116 158 146 162 154 176 161 144 134 130 100 114 112 89 129 47 99 71 142 126 160 131 74 68 108 151 154 144 140

80.6 87.6 69.2 83.4 90.0 85.4 135.0 89.2 84.7 90.1 84.5 95.9 90.7 91.3 80.3 80.0 91.3 90.9 96.2 88.1 94.4 96.6 78.5 91.4 155.0 70.2 86.5 130.0 97.1 86.2 92.6 76.4 83.7 137.0 118.0 94.2 155.0 142.0 100.3 113.0 100.2 91.9 101.6 95.3 120.0 116.0 103.0 90.4 92.4 96.3 93.8 144.0 147.0 90.6 83.3 94.0 124.0 73.7 81.5 71.7 66.9 86.0 67.4 53.1 104.6 93.8 171.0 134.0 49.5 46.3 67.2 88.9 128.0 118.0 133.0

8.75 7.83 8.55 8.03 8.96 8.46 7.90 6.75 7.05 7.76 7.53 7.29 7.39 8.45 7.85 7.26 7.91 7.61 7.92 8.99 8.40 8.21 7.17 8.37

7.60

8.43 8.64 8.25 8.27 8.42 8.45 7.38 7.78 8.00 7.36

7.36 7.14 8.24 6.17 7.18 6.08 6.34 6.19 4.59 6.96 6.84

4.11 3.65 6.51 7.39

0.446 0.466 0.467 0.487 0.494 0.443 0.495 0.418 0.439 0.493 0.471 0.549 0.466 0.521 0.507 0.451 0.464 0.335 0.299 0.448 0.398 0.402 0.356 0.424

0.428

0.453 0.472 0.442 0.455 0.446 0.507 0.410 0.433 0.429 0.421

0.405 0.400 0.438 0.339 0.406 0.327 0.358 0.351 0.259 0.388 0.374

0.221 0.221 0.361 0.438

Cr/Al

Ti/Al

V/Cr

V/(V + Ni)

Zr/Al

41.60 32.80 24.92 30.30 36.82 38.20

0.062 0.065 0.053 0.058 0.054 0.062 0.051 0.060 0.055 0.061 0.055 0.052 0.063 0.062 0.062 0.064 0.063 0.075 0.063 0.062 0.065 0.062 0.059

32.75 14.71

0.044 0.038

28.22 24.57 25.04 29.08 34.55

0.050 0.047 0.049 0.050 0.051

46.18

0.056

36.84

0.054

36.69 31.47 24.99 39.20

0.055 0.054 0.055 0.053

43.99

0.060

46.33 44.02 38.26 34.32

0.056 0.056 0.054 0.057

46.96 38.04 43.01

0.055 0.056 0.053

100.97 85.31 86.38 106.48

0.055 0.057 0.054 0.056

84.68 57.89 84.73 71.51

0.057 0.056 0.056 0.055

58.86 71.32 50.94 74.60

0.054 0.061 0.055 0.059

0.48 0.53 0.45 0.46 0.43 0.49 0.49 0.40 0.48 0.40 0.43 0.43 0.42 0.45 0.47 0.45 0.46 0.42 0.42 0.45 0.46 0.46 0.38 0.45 0.46 0.52 0.50 0.54 0.47 0.45 0.45 0.46 0.43 0.46 0.46 0.46 0.50 0.47 0.44 0.44 0.44 0.46 0.47 0.45 0.41 0.48 0.57 0.43 0.41 0.45 0.46 0.51 0.45 0.44 0.44 0.38 0.38 0.34 0.31 0.29 0.33 0.27 0.30 0.28 0.34 0.31 0.34 0.33 0.29 0.32 0.31 0.34 0.37 0.34 0.41

10.00 11.60 9.03 10.50 11.06 11.30

35.30 38.20 37.24 39.70 32.33 33.19 38.30 36.40 36.80 42.00 47.80 41.80 46.00 35.40 47.90 75.20 35.50

0.56 0.74 0.71 0.70 0.50 0.63 0.59 0.52 0.59 0.52 0.60 0.57 0.53 0.61 0.67 0.64 0.57 0.48 0.51 0.49 0.63 0.54 0.28 0.67 0.41 0.48 0.70 0.55 0.63 0.60 0.59 0.60 0.52 0.36 0.50 0.50 0.45 0.41 0.51 0.41 0.52 0.56 0.76 0.48 0.40 0.51 0.64 0.46 0.43 0.53 0.61 0.63 0.38 0.42 0.49 0.37 0.25 0.18 0.18 0.17 0.19 0.22 0.19 0.27 0.24 0.26 0.19 0.18 0.31 0.26 0.33 0.27 0.28 0.27 0.37

10.19 10.80 10.54 10.50 10.70 10.72 11.60 11.90 11.30 11.80 12.10 13.20 11.90 11.20 12.30 10.80 11.60 9.22 10.92 10.80 10.26 11.28 10.66 10.00

12.39

11.90 11.60 11.14 12.29 11.32 13.73 12.25 11.88 12.04 12.74

12.31 11.67 11.41 11.94 11.35 11.79 10.55 10.89 11.57 15.03 13.71

12.04 12.68 10.32 12.03

(continued on next page)

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Table 1 (continued) Sample number

Bed number

Section

Formation

228 227 Light 227 Dark 218 223 222 224 190 21 63 189 19 187

64 65 65 66 67 68 69 70 71 72 73 74 74

Ildjernet Ildjernet Ildjernet Ildjernet Ildjernet Ildjernet Ildjernet Fornebu Nakkholmen Nakkholmen Fornebu Nakkholmen Fornebu

Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Frognerkilen Nakkholmen Nakkholmen

Al (%) 4.07 6.81 7.67 6.50 7.50 8.22 7.73 7.40 7.62 7.91

Cr (ppm)

Ni (ppm)

459.0 143.5 289.2 216.7 524.0 222.5 282.0 242.1 330.4 234.8 204.1 242.2 227.3

201.0 117.8 186.2 141.2 182.0 139.1 148.0 148.7 191.2 148.9 133.8 138.2 140.2

Ti (%) 0.227 0.400 0.406 0.391 0.431 0.456 0.421 0.411 0.426 0.443

V(ppm)

Zr (ppm)

127 67 112 105 138 107 128 132 171 151 126 143 130

138.0 49.5 91.3 125.6 144.0 84.8 137.0 94.7 106.1 101.8 97.7 96.2 99.2

Cr/Al

Ti/Al

35.26 42.47 28.25

0.056 0.059 0.053

34.23

0.060

32.28 40.19 30.38 27.58 31.78 28.74

0.057 0.055 0.054 0.056 0.056 0.056

V/Cr

V/(V + Ni)

0.28 0.47 0.39 0.48 0.26 0.48 0.45 0.55 0.52 0.64 0.62 0.59 0.57

0.39 0.36 0.38 0.43 0.43 0.43 0.46 0.47 0.47 0.50 0.48 0.51 0.48

Zr/Al 12.16 13.41 16.38 13.05 12.63 12.91 13.17 13.20 12.62 12.54

Biostratigraphic correlation between the investigated localities is according to Hansen (2008).

structures. The dominant trace fossils are Planolites ichnosp., Chondrites ichnosp. and a large-sized Thalassinoides ichnosp. (Fig. 2A–C). Planolites ichnosp., with a diameter of about 2–3 mm, is abundant throughout the sequence and often pyritized. Chondrites ichnosp. is somewhat less common and mostly seen by colour changes in the sediment. It appears to be most common near the formation boundaries and in the upper part of the Arnestad Formation. The trace fossil has not been observed in the interval from the basal bed of the Grefsen K-bentonite complex to just below the lower Grimstorp Kbentonite bed, and in the middle part of the Frognerkilen Formation. Most specimens of Chondrites have a tube diameter of about 1– 1.5 mm, but in the lower part of the Frognerkilen Formation, and in the upper metres of the Frognerkilen Formation at Nakkholmen, they are large with tubes up to 8 mm in diameter. Thalassinoides ichnosp. occurs in both the Vollen and Frognerkilen formations dominated by carbonate mudstones, but is especially common in the lower half of the Frognerkilen Formation. The ichnospecies shows nearly no stratigraphical overlap with the Chondrites ichnosp. The specimens are generally somewhat modified by early diagenetic growth of carbonate concretions. Thalassinoides ichnosp. has a burrow diameter of 4–5 cm. 4.2. Body fossils The shelly soft-bottom fauna shows an overall high diversity with a mixture of well-preserved specimens and fragments of brachiopods, bryozoans, conchostracids, conulata, echinoderms, graptolites, machaerids, molluscs, ostracods and trilobites. However, just below and above the Kinnekulle K-bentonite, there are few

fossils and the meagre fauna is dominated by pelagic graptolites. Generally, brachiopods are the dominating macrofossils constituting about one-third of the total fossil fauna. Five brachiopod assemblages are recognized in the succession. These are, in ascending order, the Grorudia–Septorthis, Chonetoidea–Osloella, Grorudia–Onniella, Chonetoidea–Onniella and Cremnorthis–lingulid assemblages. The stratigraphic range of the brachiopod genera and assemblages in the studied area is illustrated in Fig. 3. The major appearances of new taxa take place in the one-metre interval from just below the upper Grimstorp K-bentonite bed and upward and, where exposed at locality Vollen 3, again from about 7.5 to 9 m above the Grimstorp K-bentonites (the two levels marked by arrows on Fig. 3). In the uppermost part of the Arnestad Formation and the Frognerkilen Formation, several genera reappear which otherwise were only known from the lowermost part of the Arnestad Formation or older strata in the area. The large-sized Kiaeromena (K.) kjerulfi (up to about 7 cm wide) is characteristic of the Frognerkilen Formation, though the occurrence is very sparse. Colonies of the bryozoan genus Diplotrypa are small, about 2– 5 mm, in the investigated succession, except in the lower part of the Frognerkilen Formation, where they become large (up to 11 cm). The large bryozoans have growth stages revealing that reworking and partial burial took place several times during the life of the colony. 4.3. Articulated to disarticulated brachiopod valve ratio Changes in the articulated to disarticulated ratio of the brachiopod valves are illustrated in Fig. 4. Generally, less than ten percent of the

Fig. 2. (A) Planolites ichnosp. within the lower part of the Arnestad Formation on the island of Persteilene. (B) Chondrites ichnosp. within the uppermost part of the Arnestad Formation in the section on the island Nakkholmen. (C) Thalassinoides ichnosp. within the lower part of the Frognerkilen Formation in the section on the southeastern spit of Nakkholmen. Scale bars are 1 cm; the hammer is 28 cm long.

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Fig. 3. Composite range chart for the brachiopod genera and recognized assemblages. Dashed lines indicate no record of the genera in the investigated profiles, but that they have been recorded in older and/or younger strata (e.g. Spjeldnæs, 1957; Harper, 1986; Webb, 1990; Hansen, 2008). Kukr. Stage — Kukruse Stage; N. Fm — Nakkholmen Formation; V. Fm — Vollen Formation; G — Grefsen K-bentonite complex; S — Sinsen K-bentonites; K — Kinnekulle K-bentonite; Gr — Grimstorp K-bentonite. The two gaps indicate stratigraphic gaps of uncertain thickness in the composite log. Arrows indicate the two levels at which the major first appearances of new brachiopod taxa take place. Assemblages: G–S — Grorudia– Septorthis; C–l — Chonetoidea–Osloella; G–O — Grorudia–Onniella; C–O — Chonetoidea–Onniella; Cr–l — Cremnorthis–lingulid.

specimens are articulated, but some variations are seen. In the lower part, there is a distinct increase from about 0 and up to 25% articulated valves followed by an excursion towards low values. In the middle part

of the Arnestad Formation, this trend is succeeded by a new increase, reaching a maximum at the Kinnekulle K-bentonite. Above that, at the Grimstorp K-bentonites, the ratio has returned to about 1–3%, from

274 J. Hansen et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 271 (2009) 268–278 Fig. 4. The ratio between articulated and disarticulated valves, V/(V + Ni) ratio, Cr content, Ti/Al, Zr/Al and Cr/Al ratios of the shaly mudstones, sea-level curves according to Kaljo et al. (2004) and Nielsen (2004). Kukr. Stage — Kukruse Stage.

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where a weak third positive excursion is initiated. This excursion never reaches values above 7% before the occurrence of an abrupt drop to values around 0–1% at the Arnestad–Frognerkilen formation boundary. Up through the remaining part of the succession, the mean ratio first slowly, then markedly, increases again. 4.4. Trace and major elements For the lower part of the studied sequence, the V/(V + Ni) ratio shifts from around 0.5 to a fairly constant level of around 0.45 up to the stratigraphic gap below the Kinnekulle K-bentonite (Fig. 4). The ratio increases from about 0.47 to 0.54 in the interval from just below the Kinnekulle K-bentonite and up to slightly below the Grimstorp Kbentonites. At the Grimstorp K-bentonites, the V/(V + Ni) ratio drops to a level of about 0.45. Barring a few minor oscillations, the ratio remains at that level until just below the limestone-dominated Frognerkilen Formation. A major negative excursion takes place at the formation boundary, reaching peak values of about 0.28 in the lower part of the Frognerkilen Formation. This major excursion, which is seen in all localities examined, is succeeded by a moderately steep return to high values of about 0.5. Samples from the shales bounding the three upper K-bentonites show no consistent changes in V/(V + Ni) ratios. The oscillations in the V/(V + Ni) ratio are identical at all localities examined, though the peak values differ slightly (Fig. 4). With the exception of the two lowermost samples from the Fornebu locality, which show extremely high ratios, all the localities exhibit nearly the same ratios at a given stratigraphic level. The variations of the V/(V + Ni) ratio are identical to those of the V/Cr ratio (Table 1), although the V/Cr ratio is strongly influenced by an ophiolitic component (see later). The Ti/Al and Zr/Al ratios both show only minor oscillations up through the investigated succession (Fig. 4), though both ratios exhibit a short-lived negative shift at the Kinnekulle K-bentonite. At the base of the Frognerkilen Formation, only the Zr/Al ratio takes a positive excursion followed by a rapid return to moderately high values. The Cr content and Cr/Al ratio both show a major positive excursion at the base of the Frognerkilen Formation, succeeded by a gradual return to previous values (Fig. 4). 5. Discussion Trace element ratios such as the whole rock V/(V + Ni) ratio have proved effective for distinguishing variations in the oxygen content of bottom waters (Hatch and Leventhal, 1992; Arthur and Sageman, 1994; Wignall, 1994; Hoffman et al., 1998; Rimmer et al., 2004). The trace elements Ni and V are typically enriched in dark fine-grained marine sediments deposited under oxygen-depleted conditions. In such conditions, V is reduced to reactive species of lower valence and enriched by the adsorption and formation of organic complexes. The enrichment of Ni (unchanged valence) is due to the formation of insoluble sulfides in the presence of dissolved hydrogen sulfide (Wehrli and Stumm, 1989; Breit and Wanty, 1991; Wanty and Goldhaber, 1992; Calvert and Pedersen, 1993). A V/(V + Ni) ratio higher than 0.84 indicates the presence of a strongly stratified water column with sulfidic bottom water. A ratio between 0.54 and 0.72 indicates anoxic bottom-water conditions, whereas an intermediate ratio between 0.46 and 0.60 indicates dysoxic conditions. Oxic conditions are indicated by a ratio less than 0.46, according to Hatch and Leventhal (1992). We utilize the value 0.57 to discriminate between dysoxic and anoxic bottom-water conditions, though care is taken in interpretation. The V/(V + Ni) ratio indicates that most of the Arnestad Formation was deposited in an environment transitional between oxic and dysoxic bottom-water conditions (Fig. 4). In support of this is a low-diverse ichnofauna consisting of Planolites ichnosp. and Chondrites ichnosp. throughout the Arnestad Formation, which indicates relatively low oxygen levels during deposition, but sufficient oxygen

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levels in the bottom waters to sustain life (Bromley and Ekdale, 1984; Savrda and Bottjer, 1986; Ekdale and Mason, 1988). Chondrites ichnosp. is consistently small in size. A narrow stratigraphical interval, including the Kinnekulle K-bentonite, is characterized by somewhat lower dysoxic conditions. Here the fauna became low-diverse with a dominance of pelagic graptolites prior to the deposition of the thick Kinnekulle K-bentonite, though the deposits remained bioturbated. The V/(V + Ni) ratio indicates a transition into oxic conditions in the upper part of the Arnestad Formation. This trend continues into the lower part of the Frognerkilen Formation before dysoxic conditions gradually return at the top of this formation and the overlying Nakkholmen Formation (Katian Stage). The bryozoan, Diplotrypa, also shows a gradual change from smaller to large colonies, which show clear signs of an increased level of water turbulence in the oxic lower part of the Frognerkilen Formation followed by a return to smaller colonies with the return to dysoxic conditions. It is worth mentioning here that the oxic interval coincides with where Hansen (2008) argued the depositional environment to have been above the normal storm wave base, whereas the rest was from deeper environments. The new data indicate a stratification of the water column with the more oxygen-depleted bottom water in accordance with the oceanic model for the Late Ordovician proposed by Railsback et al. (1990), and support the existing bathymetric interpretation of the area. Among the brachiopods, a Cremnorthis–lingulid assemblage characterizes the oxic interval, and is stratigraphically preceded and succeeded by a diverse Chonetoidea–Onniella assemblage. The Chonetoidea–Onniella assemblage, like the Chonetoidea–Osloella assemblage occurring here in the lower part of the studied succession, is known to have preferred offshore environments, with the Chonetoidea–Osloella assemblage (equal to the Chonetoidea–lingulid assemblage) occurring further offshore (e.g. Harper and Pickerill, 1996). Our investigation indicates that the oxygen level in the bottom water was about the same in the environment occupied by the Chonetoidea–Onniella assemblage as in the environment occupied by the Chonetoidea–Osloella assemblage. Thus, this change in brachiopod assemblage was not related to variations in the oxygen level of the bottom water. Sea-level changes are often indicated by the articulated to disarticulated ratio of brachiopod shells, as the ratio is generally higher during highstand than lowstand conditions. Differences in the strength of the hinge between different soft-bottom taxa are generally of minor importance in the brachiopod fauna of this study, and therefore differences in the ratio between articulated to disarticulated brachiopod shells have been related to post-mortem reworking and transportation which is strongly dependent on the hydrodynamic energy regime. Thus, less post-mortem reworking of the brachiopod shells and a higher ratio of articulated to disarticulated brachiopod shells will be expected from a tranquil sedimentary environment during highstand compared with skeletal material derived from an agitated environment at a shallow water depth (e.g. Brenchley and Harper, 1998). There are no observations in our material indicating differences in the strength of the hinge of these brachiopods, which could explain the variation in the ratio of articulated and disarticulated valves. By combining data on the articulated to disarticulated ratio and the V/(V + Ni) ratio, our data are closely comparable to sealevel curves proposed by Kaljo et al. (2004) and Nielsen (2004) for Baltoscandia (Fig. 4). Nielsen (2004) argued that these Baltoscandian sea-level changes were eustatic in nature. Our data indicate an overall trend from a transgressive to a highstand phase in the basal part of the succession, followed by a slow regression and then transgression again. In the upper part, there is evidence of a general regression from a highstand at the Kinnekulle K-bentonite to an absolute lowstand in the lower part of the Frognerkilen Formation, followed by a major transgression. Here, it can be noted that graptolites were generally most important constituents in the parts of the sequence which appear to be transgressive and highstand facies. This agrees with what Nielsen (2004) found in his study. The overall trend with most

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shallow-water facies within the limestone-dominated formations is also in harmony with the conclusion of Möller and Kvingan (1988) in their sedimentological study of Upper Ordovician and Lower Silurian deposits in the Oslo Region. In the present study, dysoxic conditions prevailed with only minor variations except during the maximum lowstand in the lower Katian Stage, where well-oxygenated bottomwater conditions prevailed. Two major first appearances of brachiopods took place near the Sandbian–Katian Stage boundary, after the cessation of the lower dysoxic conditions but prior to the major oxic and shallowing event (Fig. 3). These first appearances took place markedly earlier than in the East Baltic area, where the major faunal turnover of brachiopods and many other fossil groups took place in the latest Keila to the earliest Oandu Stage (Harper and Hints, 2001; Kaljo et al., 2004). However, where the East Baltic faunal turnover was mostly a replacement of the pre-existing fauna during a major lowstand event with subaerial exposure, the faunal changes in the Oslo Region were mostly through a supplement of new taxa. Only later, at the same time as the turnover in the East Baltic (e.g. Kaljo et al., 2004), did the old taxa in the Oslo Region start to disappear (Fig. 3). In eastern USA, Patzkowsky and Holland (1997) described a faunal turnover much like and possibly simultaneous with the one in East Baltica, suggesting that the extinction event was more widespread than just on Baltica. It has been suggested that the major influx of new taxa on the northwestern offshore part of the Baltoscandian Sea as well as in the shallower eastern part of the sea was in response to a regression, climatic change, plate movements and the general offshore expansion of taxa. These changes led to the migration of taxa between continents and from shallower environments on Baltica (Harper and Hints, 2001; Hints and Harper, 2003; Ainsaar et al., 2004; Hansen and Harper, 2008). Taxa are seldom restricted by water depth as such, but by other factors related to bathymetry and current systems like food, light, substrate and water chemistry (Emig, 2000; Richardson, 2000; Thomsen, 2001). In this case, neither substrate nor oxygen content in the bottom water appears to have been the triggering factor for the influx of new taxa. Most of the brachiopod genera present in the studied area just before the shallowing were also present through the regression itself when there was an abrupt increase in the oxic level. These were joined by both several older genera (e.g. Acanthambonia, Glyptorthis, Porambonites, Pseudopholidops), which had been absent through most or all of the Arnestad Formation, and some new (e.g. Kiaeromena, Protozyga, Spinilingula). The following transgressive event with a return to dysoxic environments seems to have had a stronger impact on the fauna, as many taxa, both old and new, disappeared for good in the area (Harper, 1986). The more pronounced turnover in the East Baltic, though apparently synchronic with that in the extinction event in the Oslo Region, may not be a result of the same factors. It is just as likely to be a result of subaerial exposure removing the original fauna and thereby opening for the settlement of new taxa with the return of the sea. The Ti/Al and Zr/Al ratios have been used as indicators of siliciclastic grain size, sedimentation rate and characterization of the sediment source area (e.g. Brumsack, 2006). Ti- and Zr-containing accessory minerals are reliable for tracking variations of the sediment source(s). The elements are normalized to Al because it is most associated with clay minerals and feldspars. A nearly constant Ti/Al and Zr/Al ratio of the Arnestad Formation indicates a rather homogenous terrigeneous sediment supply to the basin (Fig. 4). An abrupt change occurred for the Zr/Al ratio in the basal part of the Katian Frognerkilen Formation, subsequently followed by a gradual return to the previous values. The variable Zr/Al ratio suggests that the low relative sea level resulted in an expansion of the sediment source area and perhaps reduced sediment transport distance. In the present case, both ratios indicate that the source area(s) of the siliciclastic sediments changed little throughout the studied succession except for a minor excursion at the Kinnekulle K-bentonite and in the base of the Frognerkilen Formation. This indicates that tectonic movements did

not significantly influence the neighbouring land areas during the late Sandbian and early Katian. Sedimentary deposits enriched in Cr have in other studies been assigned to detrital material with an ophiolitic component, which is often characterized by the addition of chromite. In comparison, the influence of oxygen deficiency during deposition is negligible (Hiscott, 1984; Dill, 1986; Bock et al., 1998). The relatively high content of Cr (Fig. 4) is within or even exceeds the range of 100–700 ppm reported by Bjørlykke (1974a,b), Dypvik (1977) and Bjørlykke and Englund (1979). Bjørlykke related a high Cr content in Sandbian and lower Katian (initial increase in local stage 4a and subsequently followed by a maximum in stages 4b and 4c) to the obduction of ophiolite complexes close to Trondheim, north of the Oslo Region. These complexes, which constituted emerging island arc systems (Roberts, 2003 and references therein), comprise several ophiolite bodies including chromite ores at the transition between the Seve Nappe and Køli Nappe complexes of the upper allochthon unit of the Scandinavian Caledonides. The chromite ores, which resemble mantle portions of the ophiolite complexes, occur in the Vågå–Folldal–Røros– Brekken areas delineating the southeastern rim of the Trondheim Region (Nilsson, 1990; Nilsson et al., 1997; Nilsson, pers. com., 2008). Northeast of Trondheim, the complexes similarly have a relatively high Cr content reaching up to 3500 ppm (Furnes et al., 1992). West of the Oslo Region, ophiolite complexes of Late Ordovician or an older age also contain Cr, though only up to about 500 ppm as recorded from the Late Ordovician Solund–Stavfjord Ophiolite Complex (e.g. Furnes et al., 2003). It is a moot point whether the sediments could have been transported from the west and past the “Telemark Land” which, at that time, assumedly constituted land or a shallow marine area. Bjørlykke (1974b) reported that the sedimentary deposits in the Oslo Region contain tiny detrital chromite grains (approximately 20 μm in size). These chromite grains, which are rather resistant to mechanical abrasion during transportation, show geochemical similarities with chromite from the about 15-km2-sized Feragen ophiolite complex which is situated north of the Oslo region (Bjørlykke, 1974b; Nilsson et al., 1997). However, the ophiolite complexes northeast of Trondheim also contain a significant amount of Cr-bearing grains, mainly as chromian spinel and ferritchromite in peridotites (Furnes et al., 1988, 1992; Iyer et al., 2008). It is notable that chromite occurs disseminated throughout all of the ultramafic ophiolites, particularly the mantel peridotites (Nilsson, pers. com., 2008). However, without regard to source area, our investigations show that the enrichment of Cr is confirmed by the relatively high Cr/Al ratio (Fig. 4), coinciding with the major drop in sea level during the late Keila (earliest Katian Stage). This indicates that the obducted ophiolite complexes were subaerially exposed and subjected to weathering and erosion during the regression phase. Some of the erosional products were transported towards the Oslo Region and in this way the subaerially exposed ophiolite complexes became an additional source area of clastic sediments for the Oslo Region. During the succeeding Oandu transgression, the supply of Cr-rich sediments decreased, suggesting that the ophiolite complexes became flooded and thereby became inactive as a sediment source. The early Katian Cr-anomaly during the short-lived regression phase can thus be used as a valuable geochemical marker in the Oslo Region, but probably also in adjacent areas along the Scandinavian Caledonides. 6. Conclusion V/(V + Ni) ratios show that, disregarding sea-level changes, nearly stable dysoxic bottom-water conditions prevailed in the deeper northwestern part of the Baltoscandian Sea during the late Sandbian and early Katian. The exception is during the major shallowing in the early Katian, when oxic conditions were abruptly introduced as the depositional environment came within a normal storm wave base. This has been found to be in agreement with the low-diverse though

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abundant ichnofauna dominated by small specimens of the ichnogenera Chondrites and Planolites dominating both below and above the shallowing event. Nonetheless, the late Sandbian to earliest Katian epibenthic fauna in this area was rich and diverse and the two major immigrations of new brachiopod taxa, which took place earlier than in the East Baltic, are seen before the shallowing. This indicates that the first appearance of taxa originating both from the shallower water in the eastern part of the Baltoscandian Sea and from non-Baltic sea areas was not the result of significantly more oxic environments. Most of the dysoxic brachiopod fauna survived the abrupt environmental change and was supplemented by several genera last seen in the basal Arnestad Formation or older strata. A gradual return to dysoxic conditions occurred during the lower part of the Katian Stage, only within about 0.5 My (equivalent to the Baltic Oandu Stage). Simultaneously with this return, a significant part of the original brachiopod taxa seem to become extinct in the area together with many of the new ones. All the five brachiopod assemblages recognized in the sequence appear to respond to sea-level changes more than lithofacies. Only one, the Cremnorthis–lingulid assemblage found in the lowstand facies, may be responding to an increasing oxygen level, but could just as well appear in response to increasing water turbulence. The other assemblages, including the two more deep-water assemblages also known from other parts of the world, the Chonetoidea–Osloella and the Chonetoidea–Onniella assemblages, did not replace each other in response to the changing oxygen level as this was nearly constant. A short-term drop in sea level resulted in subaerial exposure, weathering and partial erosion of the ophiolitic complexes north but less likely west of the Oslo Region. The short-lived subaerial exposure and partial erosion of the ophiolite complexes north of the Oslo Region can thus be traced as a Cr-anomaly in the basal Katian Stage. The Cranomaly is thus a potential geochemical marker horizon in southeastern Norway. The depositional conditions were consistent within the outer dark grey mudstone dominated facies belt of the Oslo Region and correspond well to sea-level curves from the central Baltica. Acknowledgements This work was kindly supported by the University of Tromsø, Norway. We thank David Harper, Jan Kresten Nielsen, Elsebeth Thomsen, Arne Thorshøj Nielsen, Dimitri Kaljo and Lars Petter Nilsson for discussing certain topics of the work. Both reviewers, Peter Sheehan and one unknown, are acknowledged for their helpful comments which improved the manuscript. Anne Gundersen is thanked for the drawings and Odile Wallerath for helping in the laboratory. References Ainsaar, L., Meidla, T., 2001. Facies and stratigraphy of the middle Caradoc mixed siliciclastic–carbonate sediments in eastern Baltoscandia. Proceedings of the Estonian Academy of Sciences, Geology 50, 5–23. Ainsaar, L., Meidla, T., Martma, T., 2004. The Middle Caradoc Facies and Faunal Turnover in the Late Ordovician Baltoscandian palaeobasin. Palaeogeography, Palaeoclimatology, Palaeoecology 210, 119–133. Arthur, M.A., Sageman, B.B., 1994. Marine black shales: depositional mechanisms and environments of ancient deposits. Annual Review of Earth and Planetary Sciences 22, 499–551. Bentor, Y.K., 1980. Marine phosphorites-geochemistry, occurrence, genesis: a symposium. Special publication, 29. Society of economic Paleontologists and Mineralogists, Tulsa. 249 pp. Bergström, S.M., Huff, W.D., Kolata, D.R., Bauert, H., 1995. Nomenclature, stratigraphy, chemical fingerprinting, and areal distribution of some Middle Ordovician Kbentonites in Baltoscandia. Geologiska Föreningens i Stockholm Förhandlingar 117, 1–13. Bjørlykke, K., 1974a. Depositional history and geochemical composition of Lower Palaeozoic epicontinental sediments from the Oslo Region. Norges geologiske Undersøkelse 305, 1–81. Björlykke, K., 1974b. Geochemical and mineralogical influence of Ordovician Island Arcs on epicontinental clastic sedimentation. A study of Lower Palaeozoic sedimentation in the Oslo Region, Norway. Sedimentology 21, 251–272.

277

Bjørlykke, K., Englund, J.O., 1979. Geochemical response to Upper Precambrian rift basin sedimentation and Lower Palaeozoic epicontinental sedimentation in south Norway. Chemical Geology 27, 271–295. Bock, B., McLennan, S.M., Hanson, G.N., 1998. Geochemistry and provenance of the Middle Ordovician Austin Glen Member (Normanskill Formation) and the Taconian Orogeny in New England. Sedimentology 45, 635–655. Bockelie, J.F., 1978. The Oslo Region during the Early Palaeozoic. In: Ramberg, I.B., Neumann, E.-R. (Eds.), Tectonics and geophysics of continental rifts. D. Riedel Publishing Company, Holland, pp. 195–202. Breit, G.N., Wanty, R.B., 1991. Vanadium accumulation in carbonaceous rocks: a review of geochemical controls during deposition and diagenesis. Chemical Geology 91, 83–97. Brenchley, P.J., Harper, D.A.T., 1998. Palaeoecology: ecosystems, environments and evolution. Chapman & Hall, London, United Kingdom. 402 pp. Brenchley, P.J., Newall, G., Stanistreet, I.G., 1979. A storm surge origin for sandstone beds in an epicontinental platform sequence, Ordovician, Norway. Sedimentary Geology 22, 185–217. Bromley, R.G., Ekdale, A.A., 1984. Chondrites: a trace fossil indicator of anoxia in sediments. Science 224 (4651), 872–874. Brumsack, H.-J., 2006. The trace metal content of recent organic carbon-rich sediments: Implications for Cretaceous black shale formation. Palaeogeography, Palaeoclimatology, Palaeoecology 232, 344–361. Bruton, D., Harper, D.A.T., 1988. Arenig-Llandovery stratigraphy and faunas across the Scandinavian Caledonides. In: Harris, A.L., Fettes, D.J. (Eds.), The CaledonianAppalachian Orogen. Geological Society Special Publication, vol. 38, pp. 247–268. Bruton, D.L., Owen, A.W., 1982. The Ordovician of Norway. In: Bruton, D.L., Williams, S.H. (Eds.), Field Excursion Guide. IV International Symposium on the Ordovician System. Paleontological contributions from the University of Oslo, vol. 279, pp. 122–131. Bruton, D.L.A., Lindström, M., Owen, A.W., 1985. The Ordovician of Scandinavia. In: Gee, D.G., Sturt, B.A. (Eds.), The Caledonide Orogen – Scandinavia and related areas. John Wiley & Sons Ltd, Chichester, pp. 273–282. Brøgger, W.C., 1887. Geologisk kart over øerne ved Kristiana. Nyt Magazin for Naturvidenskaberne 31 36 p., map. Calvert, S.E., Pedersen, T.F., 1993. Geochemistry of Recent oxic and anoxic marine sediments: implications for the geological record. Marine Geology 113, 67–88. Cocks, L.R.M., Torsvik, T.H., 2004. Major terranes in the Ordovician. In: Webby, B.D., Paris, F., Droser, M.L., Percival, I.G. (Eds.), The great Ordovician biodiversification event. Columbia University Press, New York, pp. 84–93. Cocks, L.R.M., Torsvik, T.H., 2005. Baltica from the late Precambrian to mid-Palaeozoic times: The gain and loss of a terrane's identity. Earth-Science Reviews 72, 39–66. Dill, H., 1986. Metallogenesis of early Palaeozoic graptolite shales from the Graefenthal Horst (northern Bavaria—Federal Republic of Germany). Economic Geology 81, 889–903. Dunning, G.R., Pedersen, R.B., 1988. U/Pb ages of ophiolites and arc-related plutons of the Norwegian Caledonides: implications for the development of Iapetus. Contribution to Mineralogy and Petrology 98, 13–23. Dypvik, H., 1977. Mineralogical and geochemical studies of Lower Palaeozoic rocks from the Trondheim and Oslo Regions, Norway. Norsk Geologisk Tidsskrift 57, 205–241. Ekdale, A.A., Mason, T.R., 1988. Characteristic trace-fossil associations in oxygen-poor sedimentary environments. Geology 16, 720–723. Emig, C.C., 2000. Ecology of inarticulated brachiopods. In: Kaesler, R.L. (Ed.), Treatise on invertebrate paleontology, Part H (revised), Brachiopoda 1. Geological Society of America and the University of Kansas Press, Boulder, Colorado, pp. 473–495. Furnes, H., Hellevang, B., Robins, B., Gudmundsson, A., 2003. Geochemical stratigraphy of the lavas of the Solund-Stavfjord Ophiolite Complex, W. Norway, and magmachamber dynamics. Bulletin of Volcanology 65, 441–457. Furnes, H., Pedersen, R.B., Hertogen, J., Albrektsen, B.A., 1992. Magma development of the Leka Ophiolite Complex, central Norwegian Caledonides. Lithos 27, 259–277. Furnes, H., Pedersen, R.B., Stillman, C.J., 1988. The Leka Ophiolite Complex, central Norwegian Caledonides: field characteristics and geotectonic significance. Journal of the Geological Society 145, 401–412. Hagemann, F., Spjeldnæs, N., 1955. The Middle Ordovician of the Oslo Region, Norway, 6. Notes on bentonites (K-bentonites) from the Oslo-Asker District. Norsk Geologisk Tidsskrift 35, 29–52. Hansen, J., 2008. Upper Ordovician brachiopods from the Arnestad and Frognerkilen formations in the Oslo-Asker district, Norway. Palaeontos 13 99 pp. Hansen, J., Harper, D.A.T., 2006. Brachiopod bio- and ecostratigraphy in the lower part of the Arnestad Formation (Upper Ordovician), Oslo Region, Norway. Norwegian Journal of Geology 86 (4), 403–413. Hansen, J., Harper, D.A.T., 2008. The late Sandbian — earliest Katian (Ordovician) brachiopod immigration and its influence on the brachiopod fauna in the Oslo Region, Norway. Lethaia 41, 25–35. Harper, D.A.T., 1986. Distributional trends within the Ordovician brachiopod faunas of the Oslo Region, south Norway. In: Racheboeuf, P.R., Emig, C.C. (Eds.), Les Brachiopodes fossiles et actuels. Biostratigraphie du Paléozoique, vol. 4, pp. 465–475. Harper, D.A.T., Hints, L., 2001. Distribution and diversity of Ordovician articulated brachiopods in the East Baltic. In: Brunton, C.H.C., Cocks, L.R.M., Long, S.L. (Eds.), Brachiopods, past and present. . Systematics Association Special, Vol. 63. Taylor & Francis, London, pp. 315–326. Harper, D.A.T., Pickerill, R.K., 1996. Mid Ordovician commensal relationships between articulate brachiopods and a trepostome bryozoan from eastern Canada. Atlantic Geology 32, 181–187. Hatch, J.R., Leventhal, J.S., 1992. Relationship between inferred redox potential of the depositional environment and geochemistry of the Upper Pennsylvanian

278

J. Hansen et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 271 (2009) 268–278

(Missourian) Stark Shale Member of the Dennis Limestone, Wabaunsee County, Kansas, U.S.A. Chemical Geology 99, 65–82. Henningsmoen, G., 1953. The Middle Ordovician of the Oslo Region, Norway 4. Ostracoda. Norsk Geologisk Tidsskrift 32, 35–56. Hints, L., Harper, D.A.T., 2003. Review of the Ordovician rhynchonelliformean Brachiopoda of the East Baltic: their distribution and biofacies. Bulletin of the Geological Society of Denmark 50, 29–43. Hints, L., Meidla, T., Nõlvak, J., Sarv, L., 1989. Some specific features of the Late Ordovician evolution in the Baltic basin. Eesti NSV Teaduste Akadeemia toimetised, Geoloogia 38, 83–87. Hiscott, R.N., 1984. Ophiolitic source rocks for Taconic-age flysch: trace-element evidence. Geological Society of America Bulletin 95, 1261–1267. Hoffman, D.L., Algeo, T.J., Maynard, J.B., Joachimski, M.M., Hower, J., Jaminski, J., 1998. Regional and stratigraphical variation in Bottomwater anoxia in the offshore Core Shales of Upper Pennsylvanian Cyclothems from the Eastern Midcontinent Shelf (Kansas), U.S.A. In: Schieber, J., Zimmerle, W., Sethi, O. (Eds.), Shales and Mudstones I. E. Schweizerbart'sche Verlagsbuchhandlung (Nägele u. Obermiller), Stuttgart, pp. 243–269. Iyer, K., Austrheim, H., John, T., Jamtveit, J.B., 2008. Serpentinization of the oceanic lithosphere and some geochemical consequences: constraints from the Leka Ophiolite Complex, Norway. Chemical Geology 249, 66–90. Jaanusson, V., 1976. Faunal dynamics in the Middle Ordovician (Viruan) of BaltoScandia. In: Bassett, M.G. (Ed.), The Ordovician System. Proceedings of Palaeontological Association Symposium, Birmingham, Sept. 1974, pp. 301–326. Jaanusson, V., Bergström, S.M., 1980. Middle Ordovician faunal spatial differentiation in Baltoscandia and the Appalachians. Alcheringa 4, 89–110. Kaljo, D., Hints, L., Martma, T., Nõlvak, J., Oraspõld, A., 2004. Late Ordovician carbon isotope trend in Estonia, its significance in stratigraphy and environmental analysis. Palaeogeography, Palaeoclimatology, Palaeoecology 210, 165–185. Kaljo, D., Nõlvak, J., Uutela, A., 1996. More about Ordovician microfossil diversity patterns in the Rapla section, northern Estonia. Proceedings of the Estonian Academy of Sciences, Geology 45, 131–148. Kennedy, W.J., Garrison, R.E., 1975. Morphology and genesis of hardgrounds and nodular chalks in the Upper Cretaceous of Southern England. Sedimentology 22, 311–386. Leslie, S.A., Goldman, D., Young, S.A., Saltzman, M.R., Nõlvak, J., 2004. The Black Knob Ridge section, southeastern Oklahoma, USA: a possible global stratotype-section and point (GSSP) for the base of the Diplacanthograptus caudatus biozone and the middle stage of the Upper Ordovician series. Geological Society of America Abstracts with programs 36 (5), 75. Männil, R., 1966. Evolution of the Baltic Basin during the Ordovician. Valgus, Tallinn. 200 pp. [In Russian wiht English summary]. Meidla, T., Ainsaar, L., Hints, L., Hints, O., Martma, T., Nõlvak, J., 1999. The midCaradocian biotic and isotopic event in the Ordovician of the East Baltic. Acta Universitatis Carolinae – Geologica 43, 503–506. Möller, N.K., Kvingan, K., 1988. The genesis of nodular limestones in the Ordovician and Silurian of the Oslo Region (Norway). Sedimentology 35, 405–420. Nestor, H., Einasto, R., 1997. Ordovician and Silurian carbonate sedimentation basin. In: Raukas, A., Teedumäe, A. (Eds.), Geology and Mineral Resources of Estonia. Estonian Academy, Tallinn, pp. 192–204. Nielsen, A.T., 2004. Ordovician sea level changes: a Baltoscandian perspective. In: Webby, B.D., Paris, F., Droser, M.L., Percival, I.G. (Eds.), The great Ordovician biodiversification event. Columbia University Press, pp. 84–93. Nielsen, A.T., Meidla, T., 2004. Late Ordovician sea level changes in the Oslo area and Estonia: tie-points and problems. In: Hints, O., Ainsaar, L. (Eds.), WOGOGOB-2004 conference materials. Tartu University Press, Tartu, pp. 71–72. Nilsson, L.P., 1990. Platinum-Group mineral inclusions in chromite from the Osthammeren ultramafic tectonite body, south central Norway. Mineralogy and Petrology 42, 249–263. Nilsson, L.P., Sturt, B.A., Ramsay, D.M., 1997. Ophiolitic ultramafites in the Folldal-Røros tract, and their Cr-(PGE) mineralisation. Norges geologiske undersøkelse Bulletin 433, 10–11. Nõlvak, J., Hints, O., Männik, P., 2006. Ordovician timescale in Estonia: recent developments. Proceedings of the Estonian Academy of Sciences, Geology 55, 95–108.

Owen, A.W., Bruton, D.L., Bockelie, J.F., Bockelie, T., 1990. The Ordovician successions of the Oslo Region, Norway. Norges Geologiske Undersøkelse Special Publication 4, 3–54. Patzkowsky, M.E., Holland, S.M., 1997. Patterns of turnover in Middle and Upper Ordovician brachiopods of the eastern United States: a test of coordinated stasis. Paleobiology 23 (4), 420–443. Pharaoh, T.C., Brewer, T.S., Webb, P.C., 1993. Subduction-related magmatism of late Ordovician age in eastern England. Geological Magazine 130, 647–656. Qvale, G., 1980. New Caradocian ostracods from the Oslo-Asker district, Norway. Norsk Geologisk Tidsskrift 60, 93–116. Railsback, L.B., Ackerly, S.C., Anderson, T.F., Cisne, J.L., 1990. Palaeontological and isotope evidence for warm saline deep waters in Ordovician oceans. Nature 343, 156–159. Riccardi, A., Kump, L.R., Arthur, M.A., D, 'Hondt, S., 2007. Carbon isotopic evidence for the chemocline upward excursions during the end-Permian event. Palaeogeography, Palaeoclimatology, Palaeoecology 248, 73–81. Richardson, J.R., 2000. Ecology of articulated brachiopods. In: Kaesler, R.L. (Ed.), Treatise on invertebrate paleontology, Part H (revised), Brachiopoda 1. Geological Society of America and the University of Kansas Press, Boulder, Colorado, pp. 441–462. Rimmer, S.M., Thompson, J.A., Goodnight, S.A., Robl, T.L., 2004. Multiple controls on the preservation of organic matter in Devonian-Mississippian marine black shales: geochemical and petrographic evidence. Palaeogeography, Palaeoclimatology, Palaeoecology 215, 125–154. Roberts, D., 2003. The Scandinavian Caledonides: event chronology, palaeogeographic settings and likely modern analogues. Tectonophysics 365, 283–299. Rõõmusoks, A., 1972. Stratigraphic breaks in the Ordovician of northern Estonia and the corresponding changes in the fauna. In: Mamet, B.L., Westermann, G.E.G. (Eds.), International Geological Congress, 24th Session, Canada 1972, Section 7. Paleontology Montreal, pp. 605–611. Savrda, C.E., Bottjer, D.J., 1986. Trace fossil model for reconstruction of paleooxygenation in bottom waters. Geology 14, 3–6. Skjeseth, S., 1952. On the Lower Didymograptus zone (3B) at Ringsaker, and contemporaneous deposits in Scandinavia. Norsk Geologisk Tidsskrift 30, 138–182 pl. 1-5. Spjeldnæs, N., 1957. The Middle Ordovician of the Oslo Region, Norway 8. Brachiopods of the Suborder Strophomenida. Norsk Geologisk Tidsskrift 37 (1), 1–214 14 pls. Sturesson, U., Popov, L.E., Holmer, L.E., Bassett, M.G., Felitsyn, S., Belyatsky, B., 2005. Neodymium isotopic composition of Cambrian–Ordovician biogenic apatite in the Baltoscandian Basin: implications for palaeogeographical evolution and patterns of biodiversity. Geological Magazine 142 (4), 419–439. Størmer, L., 1967. Some aspects of the Caledonian geosyncline and foreland west of the Baltic Shield. The Quarterly Journal of the Geological Society of London 123, 183–214. Thomsen, E., 2001. Brachiopoda in the Faroe Islands area. Fróðskaparrit 49, 109–126. Torsvik, T.H., Cocks, L.R.M., 2005. Norway in space and time: a Centennial cavalcade. Norwegian Journal of Geology 85, 73–86. Torsvik, T.H., Ryan, P.D., Trench, A., Harper, D.A.T., 1991. Cambrian–Ordovician paleogeography of Baltica. Geology 19, 7–10. Trotter, J.A., Williams, I.S., Barnes, C.R., Lécuyer, C., Nicoll, R.S., 2008. Did cooling oceans trigger Ordovician biodiversification? Evidence from conodont thermometry. Science 321, 550–554. Wanty, R.B., Goldhaber, M.B., 1992. Thermodynamics and kinetics of reactions involving vanadium in natural systems: accumulation of vanadium in sedimentary rocks. Geochimica et Cosmochimica Acta 56, 1471–1483. Webb, C., 1990. Faunadynamikk i Arnestadformationen (Mellem-Ordovicium) i OsloAsker distriktet, Norge. Unpublished Master thesis. Paleontological Museum, University of Oslo. 230 pp. Wehrli, B., Stumm, W., 1989. Vanadyl in natural waters: adsorption and hydrolysis promote oxygenation. Geochimica et Cosmochimica Acta 53, 69–77. Wignall, P.B., 1994. Black shales. Oxford Monographs on Geology and Geophysics 30. Oxford University Press, Oxford. 127 pp.