Glacioeustasy during the middle Eocene? Insights from the stratigraphy of the Hampshire Basin, UK

Palaeogeography, Palaeoclimatology, Palaeoecology 300 (2011) 84–100

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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

Glacioeustasy during the middle Eocene? Insights from the stratigraphy of the Hampshire Basin, UK Caroline F. Dawber a,⁎, Aradhna K. Tripati a,b,c,d, Andrew S. Gale e, Conall MacNiocaill f, Stephen P. Hesselbo f a

Department of Earth Sciences, University of Cambridge, Downing Street, Cambridge, CB1 2EQ, United Kingdom Institute of Geophysics and Planetary Physics, University of California, Los Angeles, CA 90095, USA Department of Earth and Space Sciences, University of California, Los Angeles, CA 90095, USA d Department of Atmospheric and Oceanic Sciences, University of California, Los Angeles, CA 90095, USA e School of Earth and Environmental Sciences, University of Portsmouth, Burnaby Road, Portsmouth, PO1 3QL, United Kingdom f Department of Earth Sciences, University of Oxford, Parks Road, Oxford, OX1 3PR, United Kingdom b c

a r t i c l e

i n f o

Article history: Received 20 January 2010 Received in revised form 19 November 2010 Accepted 16 December 2010 Available online 22 December 2010 Keywords: Relative sea level Hampshire Basin Barton Clay Formation Glaicoeustasy Bartonian Middle Eocene

a b s t r a c t The sedimentary strata of the Hampshire Basin constitute some of the best-preserved Palaeogene sequences worldwide, and include the traditional ‘unit stratotype’ for the Bartonian (~36.9–41.4 Ma). The Barton Clay Formation at Alum Bay on the Isle of Wight (IOW) was studied to assess the evidence for middle Eocene sea-level variation in records of grain size, sediment properties, faunal assemblage, foraminiferal diversity indices and foraminiferal stable isotopes. Sedimentary cycles of 1–10 Myr (third order) and 0.2–0.5 Myr (fourth order) duration are reported and interpreted to reflect ~ 20–60 metre variations in water depth. Additionally, an integrated magneto-bio-chemostratigraphical age model for the succession at Alum Bay is presented and new and published litho- and bio-stratigraphical markers are used to correlate additional successions. Based on this age model, it appears that during the late Lutetian and early Bartonian (~42–38 Ma), water depth variation identified within the basin was synchronous. Sedimentary and fossil evidence supports episodic uplift in the eastern part of the Hampshire Basin during the Bartonian, which at present precludes the calculation of eustatic sea-level. However, the amplitude and frequency of water-depth variations identified in the Barton Clay Formation, and correlations to published sea-level curves, are consistent with a component of these changes being glacioeustatic during the middle Eocene. There is also evidence for a large excursion (δ18O N 1‰) in the mono-specific benthic foraminiferal oxygen-isotope record (Alum Bay) ~ 39.9 Ma, which is correlated to the isotope excursion at the ‘middle Eocene climatic optimum’ previously reported in the Southern Ocean, and other localities. A contemporaneous water-depth increase of ~ 40 m at Alum Bay may indicate that a component of this ‘global’ oxygen-isotope excursion results from a reduction in continental ice storage. © 2011 Elsevier B.V. All rights reserved.

1. Introduction The transition from ‘greenhouse’ to ‘icehouse’ conditions during the early Cenozoic marks a fundamental shift in Earth's climate (e.g. Prothero and Berggren, 1992). Although it is widely thought that this transition commenced ~ 34 million years ago (Ma), coincident with the Eocene–Oligocene boundary (Coxall et al., 2005; Zachos et al., 2001, 2008), sequence stratigraphic, sedimentary and geochemical proxy reconstructions provide evidence of continental ice storage on Antarctica during parts the middle and late Eocene (Browning et al., 1996; Dawber and Tripati, accepted; Edgar et al., 2007; Kominz et al., 2008; Miller et al., 2005; Pekar et al., 2005; Tripati et al., 2005) and also in the northern hemisphere (Dawber and Tripati, accepted;

⁎ Corresponding author. Tel.: + 44 1223 333449; fax: +44 1223 333450. E-mail address: [email protected] (C.F. Dawber). 0031-0182/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.palaeo.2010.12.012

Eldrett et al., 2007; Moran et al., 2006; Stickley et al., 2009; St. John, 2008; Tripati et al., 2005, 2008). However, precise ice budget estimates for this interval are debated, with glacioeustatic variation ranging from 0 to 150 m (Bohaty et al., 2009; Burgess et al., 2008; Dawber and Tripati, accepted; Edgar et al., 2007; Miller et al., 2005; Pekar et al., 2005; Tripati et al., 2005). This glacioeustatic range could accommodate several scenarios for the early Cenozoic cryosphere, spanning ephemeral ice sheets solely on Antarctic (e.g. Browning et al., 1996) to major bipolar glaciation (e.g. Tripati et al., 2005). Quantitative, stratigraphically based estimates of sea-level change for the middle Eocene are restricted to the New Jersey Coastal Plain (NJCP; Browning et al., 1996; Kominz et al., 1998, 2008; Miller et al., 2005; Mitrovica, 2003), although the correlation of these cycles to sequences on the East Tasman Plateau and South Tasman Rise (Pekar et al., 2005; Rohl et al., 2004) supports a eustatic origin. The NJCP backstripped estimates of regional eustasy (Browning et al., 1996; Kominz et al., 1998, 2008; Miller et al., 2005) are consistent with some

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estimates of ‘relative’ or ‘apparent’ sea-level change (Pekar et al., 2002, 2005; ~ 45 m) that are based on benthic foraminiferal oxygen isotopes (δ18O), but are two to four times smaller than estimates based on Pacific Ocean carbonate and seawater oxygen-isotope reconstructions (~80–125 m; Dawber and Tripati, accepted; Tripati et al., 2005). It is argued that the change in the volume of ocean water from the growth and decay of continental ice sheets (i.e. ‘apparent’ sea-level) should be ~ 33–48% greater than eustasy (i.e. from backstripping), as eustasy accommodates hydro-isostatic loading (Katz et al., 2008; Kominz and Pekar, 2001; Pekar et al., 2002). However, when applied to the whole ocean, others debate the concept of ‘apparent’ or ‘iceequivalent’ sea level (see review in Mitrovica, 2003). Even with a correction, middle Eocene backstripped sea level change estimates from the NJCP are still substantially smaller (up to 80 m) than sea level estimates based on a Pacific open-ocean seawater δ18O reconstruction (Tripati et al., 2005). Discrepancies between these records may reflect several factors, including uncertainties associated with the methods used, and attendant assumptions. Additional middle Eocene sequence strati-

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graphic records of local and regional sea level are needed to better define the eustatic nature of sea-level variations inferred from the NJCP and to help address the mismatch with open-ocean geochemical proxy reconstructions of relative sea level. Presented here are sediment and foraminifera data from the Barton Clay Formation that outcrops in the Hampshire Basin, UK, which is used to assess the evidence for regional sea-level variations during middle Eocene. The Hampshire Basin contains well-preserved, shallow and marginal marine Palaeogene sediments that are exposed at a number of localities on the south coast of mainland England and on the Isle of Wight (Fig. 1). Also presented is a single-species benthic δ18O record for the Barton Clay Formation at Alum Bay, IOW. This record overlaps the proposed ‘middle Eocene Climatic Optimum’ (MECO) δ18O excursion previously reported in open-ocean and hemipelagic environments (Bohaty and Zachos, 2003; Bohaty et al., 2009; Jovane et al., 2007; Spofforth et al., 2010). The new record allows us to assess the extent of this apparent perturbation in a continental shelf setting and its relationship with regional climate and sea level variations in the Hampshire Basin.

Fig. 1. Geography and structural features of the Hampshire Basin, UK. a) Outcrop map of Palaeogene strata in the Anglo–Paris–Belgium Basin (modified from Curry, 1966). b) Successions of the Barton Clay Formation are highlighted on a bathymetric map of the English Channel (from Gupta et al., 2007).

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2. Geological setting The Hampshire Basin is an asymmetrical, elongated syncline filled with up to 750 m of Tertiary sediments (Insole et al., 1998), and is the most northwesterly component of the larger Anglo–Paris–Belgium Basin (Fig. 1). Palaeogene sediments were deposited during a series of marine transgressive–regressive cycles, and reflect depositional environments from the foreshore to the outer continental shelf (Plint, 1983, 1988). The regional climate reflected the location of the basin at a palaeolatitude of ~ 40°N (Smith et al., 1981), and the globally high Eocene temperatures (e.g. Pearson et al., 2001; Tripati et al., 2003). These conditions favoured moderate chemical weathering and resulted in widespread authigenic iron and calcium phases in middle Eocene strata (Bale, 1984; Gilkes, 1968). The Palaeogene climate of the Hampshire Basin is thought to have been relatively stable (Andreasson and Schmitz, 1996; Chateauneuf, 1980). Palynology and clay mineralogy support an increasingly seasonal climate in the Hampshire Basin from the middle Eocene (Gale et al., 2006; Hubbard and Boulter, 1983), with warm and relatively invariable summer temperatures (20–25 °C), but variable winter temperatures (Hubbard and Boulter, 1983). 3. Chronology and stratigraphy of the Barton Clay Formation The Barton Clay Formation outcropping at Barton-on-Sea and at Alum Bay has been identified by the International Subcommission on

Series

Stage

Chrons Polarity

Period

Planktonic foraminifera stratigraphy

Calcareous nannofossil stratigraphy

Berggren et al., (1995); Berggren & Pearson (2005)

Martini (1971)

Palaeogene Stratigraphy (ISPS) as a target for developing Bartonian chronostratigraphy. The traditional ‘unit stratotype’ succession at Barton-on-Sea on the southern coast of England (Fig. 1) is in a poor state owing to sediment slumping and the establishment of sea defences. The succession at Alum Bay on the Isle of Wight is generally well exposed. Despite extensive work on the Barton Clay Formation (e.g. Aubry, 1986; Bujak, 1980; Burton, 1933; Chandler and OF, 1961; Curry, 1937, 1942, 1964, 1965, 1966, 1976, 1978; Gale et al., 1999; Gardner et al., 1888; Murray and Wright, 1974; Prestwich, 1857; Todd, 1990), the stratigraphy of these target successions is poorly constrained (Fig. 2). Reported here is a magnetic polarity reversal stratigraphy and a study of calcareous nannofossils for the Barton Clay Formation succession at Alum Bay. Additional successions are correlated to this model on the basis of new lithostratigraphy and published biostratigraphical markers (Aubry, 1983, 1986; Bujak, 1980; Costa et al., 1976). 3.1. Biostratigraphy Calcareous nannofossils exhibit variable preservation, and no evidence of the marker species outlined by Martini (1971) was found. In a low-resolution study, Aubry (1983, 1986) was also unable to define precisely calcareous nannofossil zones in the succession at Alum Bay. The lack of marker species may reflect poor carbonate preservation and/or insufficient water depth or connectivity with the open ocean for coccolithophores to thrive.

Hampshire basin Sequences

Hampshire basin Magnetostratigraphy (Townsend & Hailwood 1985; Gale et al., 2006)

(Curry et al., 1978; Hooker 1980)

Hampshire basin Biostratigraphy (Aubry 1983; Costa & Downie 1976; Costa et al.1976; Bujak 1980; Knox et al., 1983; Knox 1983).

O5

N2, P21

O4

N1, P20

O3

P19

O2

NP 23

P18

O1

NP 22

1n

Oligocene

Rupelian

C11n 1r 2n C11r C12n

NP 24

Bouldnor Fm.

C12r C13n

C13r

35

Priabonian

C15n C15r 1n 1r 2n C16r 1n

P16

E16

E15

C16n C17n

P15

E14

Bartonian

C18n 1r

2n

C18r

Age (Ma)

C19r

P14 P13

P12

45

Eocene

Lutetian

E12

E11

P11

NP 17

C22n C22r C23n

Ypresian

1n 1r

C23r C24n2n/r

55

2n

3n

Selsey Fm. Marsh Farm Fm.

P10

E8

Earnley Fm. NP 14

P9

E7

E8

E6

P7

E5

NP 12

P6b

E4

NP 11

NP 13

Wittering Fm. Bagshot Beds

P6a

E3 E2 E1 P5

C25n

C25r C26n

P4

P4

London Clay Fm.

NP 8* NP 8

NP 16 C. intricatum

NP 15

P. arcuatum

Earnley

NP 14 P. comatum

NP 13

Wittering

NP 12 H.abbreviatum

London Clay III London Clay II

D. varielongitudum D. simile

London Clay I Oldhaven

Reading Fm.

H. variabile

R. draco

NP 10 NP 9

R. perforatum

R. porosum

Huntingbridge

NP 15

P5 Thanetian

NP 17

Barton Clay Fm.

1n 1r

C24r

Paleocene

(=upper Barton beds)

Naish Member (=middle Barton beds) Highcliffe Member (=lower Barton beds) NP 16 Huntingbridge Member

E9

C21r

50

? Barton Sand Fm.

Paleogene C21n

15n

Headon Hill Fm.

E10

C20n C20r

E13

13n

Bembridge Lmst. Fm.

NP 19-20

NP 18

C17r

40

NP 21

2n 3n

1n

?

Bracklesham Group

C10r

30

NP 11 W. meckelfel -densis -densis A.hyperacontium

NP 6

Fig. 2. Summary of the Palaeogene sequences of the Hampshire Basin and published magneto- and bio-stratigraphic constraints (Aubry, 1983; Bujak, 1980; Costa and Downie, 1976; Costa et al., 1976; Gale et al., 2006; Townsend and Hailwood, 1985).

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The occurrence of generic NP16 species (e.g. Renticulofenestra renticulata and R. umbilica) confirms a ‘middle Eocene’ age, though more precise calcareous nannofossil datums could not be constrained. Aubry (1983) tentatively located the NP16/17 boundary in the succession at Alum Bay on the basis of secondary marker species (highest occurrence of Discoaster distinctus) and lithostratigraphic correlation to the succession at Whitecliff Bay, also on the Isle of Wight. Due to subsequent changes in the stratigraphic nomenclature and lithostratigraphy, and also the uncertainty of assigning the boundary using secondary markers, the level of this datum could not be precisely defined. It is estimated that the horizon identified by Aubry (1983) as approximately the NP16/17 boundary, occurs approximately +18–22 m above the basal conglomerate (Fig. 3). This datum is used in the age model. The published Barton Clay Formation dinocyst zonation is based partly on the succession at Alum Bay (Bujak, 1980; Costa and Downie, 1976; Costa et al., 1976). The first occurrences of selected dinocysts are used as additional datums (Williams et al., 2000) for the Barton Clay Formation age model, but note that these datums are ‘best guess’ estimates based on the compilation of data from multiple localities (H. Brinkhuis, pers. comm. 2008). The apparent regional diachrony of some of the reported dinoflagellate cysts (e.g. Eldrett et al., 2004; Heilmann-Clausen and Van Simaeys, 2005; A. Houben, pers. comm. 2008; Williams et al., 2000) may result in small uncertainties in the absolute ages of these datums (± 0.1–0.3 Ma). 3.2. Magnetostratigraphy Hand specimens were collected at 1 metre intervals from the succession at Alum Bay (Fig. 3), along with the strike direction to

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make a bedding correction. The natural remanent magnetism (NRM) was measured on ~ 75 samples of fine sand to clay using a 2G magnetometer at the University of Oxford. The NRM was measured sequentially after thermal demagnetization steps in a field free room. Samples were systematically rotated within the furnace to minimize potential laboratory acquired magnetism. Samples were thermally demagnetized at 50 °C intervals between 100° and 450 °C, and then at 30 °C intervals between 450 and 580 °C. Of the ~ 75 samples analysed, about 25% yielded comprehensible demagnetisation histories. Those that yielded weak magnetisations or acquired re-magnetisation in the furnace were excluded. A number of samples exhibited clear demagnetisation histories, from which the characteristic remanent magnetism (ChRM) could be confidently inferred. These samples were used as tie points for the interpreted polarities. All other samples had statistically significant mean angular deviation (MAD) values b 15. The magnetic polarity reversal history for the Barton Clay Formation at Alum Bay is summarised in Fig. 4. An extensive normal polarity magnetozone in the upper part of the succession at Alum Bay (+35 m to + 72 m above the base, Fig. 4) is inferred to be Chron C18n, the longest normal polarity magnetozone during the middle Eocene. Chron C18n occurs exclusively within calcareous nannofossil zone NP17 (Berggren and Pearson, 2005). The occurrence of the secondary NP17 zonal marker Sphenolithus obtusus approximately + 25 m above the basal conglomerate at Alum Bay (Aubry, 1983) supports the interpretation of this normal magnetozone as Chron C18n. The absence of reverse magnetic polarities in the upper Barton Clay Formation at Alum Bay likely indicates an upper age limit of ~ 39.6 Ma (top of Chron C18n.2n). Although sedimentation rates vary throughout the succession (Fig. 5), it is unlikely that the sampling

Fig. 3. Sedimentary log of the Barton Clay Formation succession at Alum Bay, Isle of Wight. The alphabetical divisions of Burton (1933) employed in previous studies are inferred based on lithostratigraphy. Arrows denote the position important lithostratigraphical markers, the basal conglomerate and stairs pebble bed.

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Biostratigraphy Dinoflagellate Cysts

-Martini, 1971 -Aubry, 1983

-Bujak et al., 1980 -Williams et al., 2000

BAR-3

Calcareous Nannofossils

Sedimentary log

Depth (m)

Chronology

Palaeomagnetic Inclination Magntozones and Chrons -90

-60

-30

0

30

60

Townsend and Hailwood 90 (1985)

Model 1, this study

Model 2, this study

C18n

C18n

70 G

60 R.porosum

G

50

BAR-2

NP 17

Bartonian

G

40

G S.obtusus

G

BAR-1

Middle Eocene

40.0 Ma

40.13

30

G G

N.minutus

40.4 Ma

G

R.reticulata

20

D.distinctus

C18r

C18r

R.draco

G

40.5 Ma

S.furcatolithoides

G

GIN

NP 16

Lutetian

G

10

G

C19n

C19r G

41.26 41.52 42.54

C20n

C19n

C20n

C20n

0 --- base of Barton Clay ---

Formation

Fig. 4. Palaeomagnetic reversal history of the Barton Clay Formation succession at Alum Bay, shown on the Berggren and Pearson (2005) global polarity time scale. Normal polarities are indicated with black boxes and periods of reverse polarity by white boxes. Upwards and downwards pointing arrows denote first and last occurrence bio-stratigraphic datums respectively.

Age (Ma), Berggren & Pearson GPTS (2005) 39.0

40.0

41.0

42.0

43.0

Magnetic polarity reversa ldatums, this study Calcareous nannofossil datums, Aubry (1983)

70

Dinocyst datums, Bujak et al., (1980); Williams et al., (2000)

60

Sedimentation rates ~2 m Ma-1

D2

Depth (m)

50 ~15 m Ma-1

40 ~188 m Ma-1

30

20

~39 m Ma-1

M4

C1

D1

10 M3

0

M2 M1

Fig. 5. Depth vs. age plot for the Barton Clay Formation succession at Alum Bay. Sedimentation rates are calculated using a linear interpolation between magnetic- and biostratigraphic datums.

resolution is too coarse to alias the transient Chron C18n1r reversal (b 100 kyrs, Berggren and Pearson, 2005). Accounting for this potential uncertainty, the youngest feasible estimate of the upper Barton Clay Formation sampled in this study is 38 Ma (top of Chron C18n.1n). In the absence of a lithological break between the upper (Chron C18n) and middle Barton Clay Formation, the underlying reverse magnetozone (+ 13 m through + 31 m above the base, Fig. 4) is inferred to be Chron C18r (~ 41.3–40.1 Ma). This interpretation is consistent with biostratigraphic constraints that place the NP16/17 boundary (within Chron C18r) approximately + 18–22 m above the basal conglomerate (Aubry, 1983). The lowermost 5.5 m of the Barton Clay Formation record a normal polarity, but are overlain by two successive samples that yield opposite polarities (+ 6.5 m reverse, + 7.5 m normal, Fig. 4). The primary magnetisation vectors of these two samples are statistically significant (MAD values of 11 and 4.3 respectively) and, with no evidence for a lithological break in the lower Barton Clay Formation, these magnetozones could be inferred to be Chrons C19r and C19n respectively (~ 42.5–41.3 Ma, Berggren and Pearson, 2005). This interpretation implies that sedimentation rates during Chron C19 are substantially lower than during Chron C18 and the early Palaeogene (Fig. 5; Aubry, 1986; Townsend and Hailwood, 1985). These low rates are also anomalous for a shallow shelf environment (cf. fig. 10.3 Einsele, 2000) and may indicate the presence of a prolonged hiatus, or a series of shorter hiatuses, although no obvious lithological breaks were observed in the outcrop.

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In light of the apparently anomalous sedimentation rates inferred for Chron C19, two age models are presented for the lower Barton Clay Formation (Fig. 4). The first includes all recorded polarities and attributes the normal polarity magnetozone that characterises the lowermost 5.5 m of the Barton Clay Formation to Chron C20n. The second model attempts to reconcile the anomalous sedimentation rates and disregards the normal polarity recorded at + 7.5 m, attributing the lower-most normal polarity magnetozone to Chron C19n. In this model, Chron 19r and potentially parts of Chron 20n and Chron 19n represent periods of non-deposition and/or subsequent erosion of this material. The base of the Barton Clay Formation likely represents a significant disconformity. At Alum Bay, the boundary with the underlying Bracklesham Group is delineated by a thin, conglomerate bed with an erosional base (Fig. 3). This bed is present in other successions to the west, but is absent in the eastern part of the basin (Gale et al., 1999; Huggett and Gale, 1997; Todd, 1990). Whilst this bed may delineate the palaeoshoreline, Gale et al. (1999) cite the Bracklesham/Barton Clay facies succession at Whitecliff Bay, IOW, and the presence of Thalassinoides burrows into the underlying Selsey Formation, as evidence for a period of non-deposition in the eastern part of the basin, possibly related to an episode of uplift on the Sandown Pericline. Todd (1990) also argues that the absence of the Studley Wood Member (part of the Selsey Formation) at all localities east of Porchfield, IOW, (Fig. 1) reflects an eastwardly increasing hiatus at this time. The Purbeck–Isle of Wight fault system (P–IOW) that extends through the Hampshire Basin consists of a number of en echelon segments (e.g. Chadwick, 1986; cf. Insole et al., 1998; Stonely, 1982), and the magnitude and history of uplift on the individual fault segments was likely variable from location to location (Plint, 1982). Thus, it is not clear whether the major disconformity identified at the base of the Barton Clay Formation in the eastern part of the basin (Gale et al., 1999; Huggett and Gale, 1997; Todd, 1990) extended into the central and western parts of the basin. However the large pebbles and cobbles, and the presence of oxidized sandy cement, probably reflects a winnowed in situ storm beach deposit, perhaps originally analogous to the eighteen mile Chesil Beach occurring today in Dorset, South England, and provides supporting evidence for a major disconformity. Both age models for the lower Barton Clay Formation are consistent with previous interpretations of magnetozones in the underlying Bracklesham Group (Fig. 4, Aubry, 1986; Townsend and Hailwood, 1985). At present it is not possible to distinguish between these age models as this part of the succession is decalcified and barren of microfossils. Alternative dating techniques are required. 3.3. Inter-basin correlations and palaeogeographical reconstruction Using lithological and biostratigraphical markers, additional successions of the Barton Clay Formation are correlated to the expanded succession at Alum Bay (Fig. 6). The first occurrence of Nummulites species provide distinct and easily identifiable (Curry, 1937) horizons that can be correlated throughout the basin. The first occurrence (FO) of Nummulites prestwichianus in the succession at Alum Bay coincides with the transgressive systems tract (cf. Section 4). Plint (1982, 1983, 1988) interpreted lateral facies variations within Bracklesham Group (Lutetian) to reflect a palaeodepth transect along an east–west trending inlet open to fully marine conditions in the east. If this palaeogeographical reconstruction is appropriate for the Bartonian, the FO of N. prestwichianus may be governed by the position of the palaeoshoreline and could be diachronous within the basin. In contrast, Gale et al. (1999), argue for open marine conditions across the basin with a more distal northern palaeoshoreline. The size distribution of clasts in the basal conglomerate of the Barton Clay Formation at Alum Bay (Fig. 3) and Hengistbury Head (Fig. 6) is consistent with the palaeogeographical reconstruction of Gale et al.

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(1999), with marine conditions open to the influence of Atlantic storms from the west. Thus, it is assumed that the palaeogeographical reconstruction of Gale et al. (1999) is most appropriate for the Bartonian, and that the FO of Nummulites species within the basin was simultaneous. On the basis of faunal evidence, the extinction of Nummulites species in the Hampshire Basin appears to be simultaneous (Curry, 1966). The basal conglomerate bed defines the base of the Barton Clay Formation in the west of the basin and is used to correlate the successions at Christchurch Bay, Hengistbury Head and Alum Bay (Fig. 6). A second distinct bed, the stairs pebble bed, is observed in the upper part of the succession at Alum Bay (~55 m above the basal bed, Fig. 3). This bed consists of a suite of ‘exotic’ pebbles, flint, and marine fossil accumulations supported in a silty to very fine sandy matrix. Its stratigraphic context, continuously under- and overlain by fine glauconitic marine sandy silts is somewhat enigmatic and it is not recognized in other successions of the Barton Clay Formation. In contrast to the basal conglomerate, which consists predominantly of flint derived from the Chalk Group, the stairs pebble bed contains additional exotic clasts that may indicate at temporary switch in sediment source during the Bartonian. The ‘exotic’ clasts likely represent material derived from the Mesozoic sediments that infill the epicontinental basins of southern England, the English Channel and Northern France. The silicified clasts are typical of the Portland Limestone Formation (Jurassic) of southern England, whilst the chalcedony may be derived from the Upper Greensand (Cretaceous), or possibly the Portland or Purbeck Groups (Jurassic), the veined quartz is similar to the Palaeozoic rocks observed in Devon (SW England) and the jasper may have been transported from Armorica, NW France (I. West, pers. comm. 2008). Today, the overburden of the middle to late Cretaceous age Chalk varies substantially through the Hampshire Basin (150–490 m; Curry and Smith, 1975; Newell, 2001), consistent with an undulating palaeokarst surface during the Palaeogene (Newell, 2001). An isochore map for the Chalk highlights an area of thinning in the central Hampshire Basin (Newell, 2001), but existing borehole successions provide no evidence to support the exposure of Early Cretaceous and Jurassic sediments within the basin during the Bartonian (Edwards and Freshney, 1987; Todd, 1990). The clasts observed in the stairs pebble bed are more likely to be sourced from outside the basin. At Blackdown and Bincombe in western Dorset (Wessex Basin), the conglomeratic fluvial facies of the middle Eocene Bracklesham Group contain a range of exotic Palaeozoic and Mesozoic clasts (Gibbard and Lewin, 2003), and are cited as the upstream equivalent to the gravels observed to the west at Wareham (Gibbard and Lewin, 2003; Plint, 1982). The compositional similarity of clasts in the Alum Bay stairs pebble bed to those observed at Wareham and Blackdown suggest some middle Eocene sediment at Alum Bay may be sourced either directly from these beds, or from lateral equivalents of these beds or from other Mesozoic deposits in the Wessex Basin. The stairs pebble bed occurs within the transgressive systems tract (cf. Section 4) and may represent a local lag deposit. This bed is matrixsupported and the absence of erosional surfaces is somewhat atypical of a transgressive lag. Although bioturbation could account for the absence of erosional surfaces, the high percentage of mud grade matrix and the exclusivity of this bed to the succession at Alum Bay are perhaps more consistent with a local debris flow. If correct, these observations suggest that there was a source of fluvially transported Mesozoic material from the Wessex Basin draining into the western and possibly central part of the Hampshire Basin during parts of the Bartonian. 4. Depositional cycles in the Barton Clay Formation: Alum Bay 4.1. Grain size and physical properties Grain size distributions were determined by dry sieving and weighing the coarse (N 63 μm fraction) at one phi intervals. The

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Fig. 6. Correlated successions of the Barton Clay Formation, Hampshire Basin. Correlations are based on three horizons: (1) the basal conglomerate, (2) the Nummulites prestwichianus horizon and (3) the N. rectus horizon. The alphabetical divisions of Burton (1933) employed in previous studies are inferred based on lithostratigraphy. Sporadic sediment slumping in the outcrop at Alum Bay introduces a ± 2 metre error on the first appearance of N. rectus.

glaucony content of the coarse fraction is estimated from the weight of the magnetic fraction, separated using a magnetic separator with a current of 1 amp, inclined at 15°. The depositional facies of the Barton Clay Formation at Alum Bay are entirely marine (Fig. 7), and with the exception of the basal conglomerate bed, predominantly reflect water depths exceeding the fair-weather wave base (i.e. N ~ 15 m, e.g. Nichols, 1998). Consequently, Bartonian sedimentary cycles in the succession at Alum Bay are not associated with erosional horizons or the juxtaposition of facies from notably different depositional environments. Sedimentary cycles with periodicities of 106 years (third-order) and 104–105 years (fourthorder) are defined on the basis of variations in grain size, sediment sorting and the percentage of glaucony pellets in the N 63 μm sediment fraction (Fig. 7). Primary variations are not observed in the colour reflectance or magnetic susceptibility, suggesting that changes in bulk mineralogy are not significantly linked to variations in water depth. A lithological summary of the sequence stratigraphic model is presented in the Supplementary information.

4.2. Glaucony abundance The presence of glaucony in siliciclastic-dominated environments has been linked to fully marine conditions and low sedimentation rates (e.g. Haq, 1991; Loutit et al., 1988). As a result, in situ glaucony abundance has been used to identify condensed horizons and to differentiate between the transgressive (TST) and highstand systems

tracts (HSTs, e.g. Amorosi, 1995; Hesselbo and Huggett, 2001; Pekar et al., 1997). Glaucony pellets are a major constituent (up to 70%) of the N 63 μm sediment fraction in the lower Barton Clay Formation at Alum Bay. The lack of alteration and weathering features and the high iron and moderate potassium content of this glaucony are cited as evidence for it being in situ (Bale, 1984). Throughout the succession, trends in glaucony abundance exhibit some similarities to records of sediment grain size and sorting, and foraminiferal diversity indices (Section 4.3) on 104–105 year timescales (Fig. 7). Over longer periods, these trends appear to be less robust. A number of studies have demonstrated that glaucony forms, in situ, in a wider range of lithofacies than previously thought, and may be influenced by a number of environmental parameters (e.g. temperature, salinity, sediment accumulation rate, water depth; Amorosi and Centineo, 1997; Hesselbo and Huggett, 2001; Huggett and Gale, 1997). The absence of a strict correlation between glaucony abundance and 106 year (third-order) sedimentary cycles may reflect the greater variability of ‘other’ environmental conditions over long time periods (Amorosi and Centineo, 1997). It is also possible that some of the glaucony may be reworked from older sediments or transported penecontemporanously within the basin. The criteria for identifying reworked glaucony have been discussed extensively (e.g. Amorosi and Centineo, 1997; Hesselbo and Huggett, 2001; Huggett and Gale, 1997; Miller et al., 1998; McCraken et al., 1996; Odin and Dodson 1982), however there are no unequivocal indicators. Furthermore, many of the pellet characteristics associated with reworked material cannot explicitly distinguish between

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Depth (m)

Magneto- Sedimentary chrons facies

Wt% >63 µm 0

20 40 60 80

Wt% magnetic (glaucony)

Grain sorting 0 0.5 1 1.5 2 2.5

0

L*

20 40 60 80 20 25 30 35 40

91

Magnetic susceptibility

Sedimentary cycles

0 5 10 15 20 25

HST 70 G

MFS

C18n

60

TS lag

G

TST

50

G

?

40

G G

30

G

HST

G

C18r

G

20

MFS TS lag

G

TST

G

C20nr C19r

G

10

G

HST

G

0

MFS

TS lag

Fig. 7. Summary of the Barton Clay Formation sediment properties at Alum Bay, IOW. Max flooding surfaces (MFSs), and low stand systems tracts (LSTs) are allocated on the basis of grain size minima and maxima respectively. Transgressive lag deposits (TSs) are inferred from unsorted sediments. Increases in glaucony abundance are interpreted as condensed horizons and define the transgressive systems tracts (TSTs). Decreases in glaucony abundance are inferred to indicate the onset of the highstand systems tracts (HSTs).

biological reworking (i.e. bioturbation), in which the glaucony is still essentially in situ, and physical reworking associated with the transportation of derived glaucony (Hesselbo and Huggett, 2001). It has been argued that when glaucony co-occurs with sand grade siliciclastic sediment, the glaucony is most likely derived (e.g. Kominz et al., 1998), however others have questioned this interpretation in deeper slope environments (Hesselbo and Huggett, 2001). The presence of reworked glaucony has implications for the sequence stratigraphical model, for example, in Oligocene New Jersey sequences, in situ glaucony occurs within the transgressive systems tract, whilst reworked glaucony is present in the highstand systems tract (Pekar et al., 1997). Although no relationship between glaucony abundance and sediment percentage in the N 63 μm fraction is observed in the Barton Clay Formation at Alum Bay, variations observed solely in glaucony abundance are not used to infer sequence stratigraphical cycles for the Barton Clay Formation at Alum Bay.

Foraminiferal diversity indices for the Barton Clay Formation at Alum Bay exhibit trends that are consistent with cycles inferred from the sediment properties (Figs. 7 and 8). Foraminiferal species richness and evenness appear to be sensitive to environmental change and are interpreted to reflect variations in water depth. The inferred increases in water depth at ~ 17 m (C18r) and 55 m (C18n, Fig. 8) coincide with fining up trends and increases in glaucony abundance. Throughout the middle Barton Clay Formation (30–50 m above the base) diversity indices fluctuate, and could indicate that the sampling resolution in this part of the succession is too low to resolve fourth order cycles or water depth variations below the threshold of those that can be identified by foraminiferal assemblages. 5. Palaeoenvironmental reconstruction and water depth variations 5.1. Foraminiferal assemblages and palaeoecology

4.3. Foraminiferal diversity indices Patterns in foraminiferal diversity have been shown to be useful for recognizing sequence boundaries and delineating systems tracts when integrated with additional palaeobathymetry data (Armentrout et al., 1991). Reduced species abundance and simple species diversity are hypothesised to mark the sequence boundary, increase through the lowstand and transgressive systems tracts, with peak values denoting the maximum flooding surface (Armentrout et al., 1991).

The modern day distribution and palaeoecology of foraminifera provides useful analogies for the reconstruction of palaeoclimate parameters including water depth. Forty-five taxonomically distinct foraminiferal tests were identified in the Barton Clay Formation at Alum Bay (full species abundance data are presented in the Supplementary information). Scanning electron micrograph (SEM) images and rarefaction are used to evaluate taphonomic and taxonomic biases.

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Margalef C. ungerianus Buzas and Gibson Species richness Dominance C. pygmeus per g sed evenness 0 1 2 3 4 5 0 0.2 0.4 0.6 0.8 1 0 10 20 30 40 50 0 0.2 0.4 0.6 0.8 1 0

Magneto- Sedimentary chrons facies

70

0.5

H 1

1.5

2 0

Fisher alpha 1 2 3 4

Sedimentary cycles 5

HST

G

C18n

60

MFS

TST

G

50

? G

40 MFS?

G G

30

G

?

G

HST C18r

G

20

TST

G G

C20nr C19r

G

10

G G

0

Fig. 8. Foraminiferal diversity indices for the Barton Clay Formation succession at Alum Bay, IOW. Trends in foraminiferal species richness and evenness are inferred to reflect sedimentary cycles. Systems tracts abbreviations as in Fig. 7.

Planktonic foraminifera are rare in the present study of the Barton Clay Formation at Alum Bay, with only one species observed. In contrast, Murray and Wright (1974) documented planktonic foraminifera at many horizons at Alum Bay, albeit in small numbers. Parts of the succession at Alum Bay are decalcified as a result of sediment exposure, which may account for the restricted planktonic assemblage observed at present. It is also possible that planktonic foraminifera may not have been able to thrive due to restricted access to the open-ocean or extreme shallow water depths. In contrast, benthic foraminifera occur in abundance and the assemblage is diverse. The majority of hyaline and porcelaneous benthic foraminifera appear ‘glassy’ and ‘pearly’ respectively, which has been suggested to reflect very good preservation (e.g. Pearson et al., 2001; Sexton et al., 2006). This inference is supported by the preservation of sub-micron crystallites of the primary calcite test (Fig. 9). The interstices of many of the benthic foraminifera are infilled with pyrite, an early diagenetic product. The oxidation of pyrite following aeration of the host sediment may have caused a reduction in pH that in turn could cause the inner part of the test to dissolve. SEM images of test interiors displayed no evidence of thinning or etching that would indicate internal dissolution (Fig. 9). The discrepancy between the exquisite benthic foraminiferal preservation and the absence of planktonic foraminifera is enigmatic, and may reflect differences in the post-depositional dissolution of benthic and planktonic foraminifera or the primary absence of planktonic foraminifera. Thus, only the benthic foraminifera are considered in this study. 5.1.1. Foraminiferal faunules The Barton Clay Formation at Alum Bay is subdivided into seven units based on foraminiferal faunules. Key species and palaeoecological interpretations based on the ecology and reported water depth distribution of extant taxa (Jenkins and Murray, 1989; Katz et al., 2003; Murray, 1971, 1991, 2006) are summarized in Table 1. Additionally, factor analysis was used to investigate associations of benthic foraminifera. Several scenarios were considered using the full and restricted foraminifera data set from Alum Bay, allowing the XLSTAT

macro to automatically and forcibly identify new factors. Only two species exhibited significant factor loadings (Cibicides pygmeus and Cibicides ungerianus). These dominant species belong to the same genus, have similar test morphology and are now extinct, so it is difficult to invoke any major differences in their ecology. It was not possible to define statistically significant biofacies using factor analysis, likely because the environmental conditions were favourable to the Cibicides species allowing them to dominate the assemblages. Species of the large benthic foraminifera (LBF) Nummulites exhibit distinct tests that could provide additional constraints on the depositional environment if in situ. However the high skeletal density makes this genus prone to post-mortem transportation. The first occurrence of N. prestwichianus in the succession at Alum Bay coincides with the TST (Figs. 7 and 8) and some specimens have broken edges that may indicate transportation. Nummulites host symbiotic algae and thrive only in waters with sufficient light and nutrient levels to allow their symbionts to photosynthesise (e.g. Beavington-Penney and Racey, 2004). The high mud (b 63 μm) content of the Barton Clay Formation succession at Alum Bay (Fig. 3) indicates a turbid shelf environment that would be detrimental to Nummulites. A reduction in sedimentation rate throughout the TST likely resulted in temporarily clearer waters and may have provided a respite for Nummulites. This rationale is supported by the reduced numbers of N. prestwichianus in the subsequent HST (Figs. 7 and 8) as turbid conditions are inferred to return. It is therefore assumed that Nummulites specimens at Alum Bay are in situ and that their test structure can be used to define more precise water depths. The flat, discoidal test of N. prestwichianus is consistent with an oligotrophic, cool water, outer-shelf (80–100 m) environment, whilst the rounder and more inflated test of N. rectus indicate a shallower (40–60 m) central shelf environment (Beavington-Penney and Racey, 2004). The deepest palaeodepth estimates constrained by the Nummulites species are consistent with the 75 m to upper shoreface amplitude cycles inferred for the lower and middle Barton Clay Formation from ostracod assemblages (Keen, 1991).

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Fig. 9. Scanning electron micrograph images of Cibicides ungerianus. Tests are composed of sub-10 micron sized primary calcite crystals that exhibit definite crystal boundaries without any signs of etching or dissolution. Note the exquisite preservation of delicate sub-micron ‘spear’ structures in the pores.

5.2. Foraminiferal morphogroups The benthic foraminiferal assemblages for the Barton Clay Formation at Alum Bay support water depth variations of 50 to 80 m (inner to outer shelf) during the late Lutetian and Bartonian (Table 1). Palaeoecology is not a stringently quantitative discipline and the use of modern relationships and the distribution of extant species may not be appropriate if significant shifts in palaeoecological niches have occurred over time. The absolute depth range of benthic foraminifera assemblages also varies from shelf-to-shelf making it difficult to precisely estimate the uncertainty associated with palaeodepths. Palaeoslope modeling using the distribution of benthic faunules at multiple sites can provide more rigorous palaeodepth estimates (e.g. Pekar and Kominz, 2001). But this was not achievable

in the Hampshire Basin owing to large sections of the Whitecliff Bay succession being decalcified. Thus, estimates of water depth variations based on foraminiferal assemblages should be viewed conservatively. Relationships between foraminiferal test morphology and habitat (e.g. Bandy, 1960) are likely to have remained more uniform and to have been less sensitive to species turnover. Constraints on key environmental parameters, including water depth are summarised in Table 2. The abundance of foraminiferal morphogroups in the Barton Clay Formation at Alum Bay supports water depth variations of 50 to 80 m (inner to outer shelf) in the lower part of the succession (e.g. 18–30 m above the basal bed; Fig. 10) and smaller amplitude variations (20–50 m) in the upper part (50–75 m above the base; Fig. 10). Although these constraints are based on a more conservative method,

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Table 1 Summary of foraminifera assemblages and palaeoecological interpretations of depositional environments in the Barton Clay Formation succession at Alum Bay. Interval

Key species

Paleoecological interpretation

Water depth estimate (m)

Basal conglomerate

No foraminifera

Beach/foreshore

Basal conglomerate to + 16 m

No foraminifera

~ 16 to 17.5 m

Trochammina sp. Cribrostomoides sp. (Documented by Murray and Wright, 1974 — not found in present study). C. pygmeus, C. ungerianus, N. prestwichianus, B. propigua, E. latidorsatum, M. affine C. pygmeus, N. rectus, C. ungerianus, B. propigua, Q. seminulum, Q. cartinata Q. reichel

Lithology inferred to reflect a winnowed in situ storm beach deposit Pyritic moulds of calcareous macrofossils and the presence of selenite/gypsum crystals indicate post-depositional dissolution. Tidal hyposaline marsh environmenta

+ 17.5 m to 26 m, straddles the N. prestwichianus horizon

+ 28 to 30 m, straddles the N. rectus horizon

+ 34–36 m

C. pygmeus,

+ 36–37 m + 39–47 m

No foraminifera C. ungerianus, C. pygmeus, B. propigua, affine, Protoelphidium

+ 54–56 m

No foraminifera

+ 57–63 m, straddling the ‘stairs pebble bed’ + 64–73 m

+ 75–78 m

C. ungerianus, B. propigua and minor occurrences of A. stelligerum C. ungerianus, B. propigua, C. pygmeus, Q. ludwigi, Q. seminulum, A. bartoniana P. spinigera C. ungerianus, B. propigua, Protoelphidium

+ 79–80 m

No foraminifera

a b c d e

Not possible

b3 m

Generic shallow marine shelf a,b,c,d. Nummulites suggest a oligotrophic, cool water, outer-shelf environmente

Central to outer shelf (50–100 m)

Generic shallow marine species together with those tolerant to more euryhaline conditions. Dominance of Cibicides indicates relatively cool shelf waters, but presence of Q. seminulum restricts to T N 10 °Cb. A seagrass or coarse sediment substrate is required for Quinqueloculina. The rounder and more inflated test of N. rectus is indicative of a central shelf environmente Assemblages dominated (72–92%) by one species, indicative of a ‘stressed environment’ (?) Post depositional dissolution horizon (?) Cibicides is the most abundant genus, but there is a marked change to the dominance of C. ungerianus over C. pygmeus. The presence of Protoelphidium (modern analogue is the brackish genus Haynesina) indicates slight hyposalinity (32–35‰)d Possible post-depositional dissolution horizon, but sharp increase in av. grain size is consistent with shoaling water depths. The depositional environment may be too shallow for benthic foraminifera Generic shallow shelf assemblagesb,c,d

Central shelf (40–60 m)

Shelf Not possible Inner to Central Shelf (20–60 m)

Not possible

Central shelf (~ 40–60 m)

The genus Quinqueloculina implies slightly euryhaline conditions, however the genus Asterigerina is thought to occur only in environments with stable salinity conditions

Inner to Central shelf (20–60 m)

The assemblage is dominated by C. ungerianus (up to 92%) indicating a limiting environmental parameter — perhaps salinity given the presence of brackish-hyposaline genus Protoelphidium Post-depositional dissolution horizon

Inner to Central shelf (20–60 m)

Not possible

Murray and Wright (1974). Murray (1971). Jenkins and Murray (1989). Murray (1991). Beavington-Penney and Racey (2004).

they are consistent with estimates of water depth derived from the palaeoecological interpretations of the foraminiferal assemblages. 5.3. Sediment grain size and percent-mud proxy Landward coarsening of sediment deposited on the continental shelf is loosely recognized as proxy for interpreting past variations in water depth. Models for reconstructing palaeobathymetric trends have been derived from investigating a number of modern wave-graded shelf environments (Dunbar and Barrett, 2005). The proxy relates the percentage of mud grade (b 63 μm) sediment to wave-induced bed shear stress, which for a constant wave climate, is governed by water depth. This approach for estimating palaeobathymetry may be overly simplistic given the uncertainty in determining past wave climates and inferring an invariant wave climate given the potential for temporal

variations in wind fields, palaeogeography, basin hypsometry and sea level (cf. Dunbar and Barrett, 2005; Dunbar et al., 2008). Nonetheless, this grain size proxy has been used to estimate glacial–interglacial water depth variations in the western Ross Sea during the Oligocene and Miocene (Dunbar et al., 2008), providing the basis for a local backstripped sea-level reconstruction (Naish et al., 2008), which appears to be consistent with other late Oligocene δ18O-to-sea level calibrations (Pekar and DeConto, 2006). Estimating palaeo-wave climate is difficult. Today, the English Channel is characterised by a moderate wave climate, with mean annual significant wave heights of ~ 1.5–1.7 m and a wave period of ~ 5–6 s (Neill et al., 2009). Model simulations have demonstrated that for the last 12,000 years the wave climate of the English Channel is less sensitive to sea level in comparison to the more open North Sea and Irish Sea (Neill et al., 2009). Although the UK was located at a

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95

Table 2 Summary of the environmental parameters governing the distribution of foraminifera morphogroups. Bays equate to 0–10 m water depth, inner shelf: 0–30 m, central shelf: 30–70 m and outer shelf: 70–120 m. Inferences on 1) water depth are based on the studies of Bandy (1960) and Severin, (1983); 2) salinity, based on Jenkins and Murray (1981) and 3) sediment substrate based on Jenkins and Murray (1981) and Murray (1991). Foraminiferal morphogroup

Water depth

Salinity

Flat discoidal Planoconvex Rounded planispiral Ovate Milliolids Elongate flattened and biconvex keeled Cylindrical

Central–outer shelf Inner–central–outer shelf Shelf Inner–central shelf, channels and bays Central–outer shelf Bays

Normal marine stenohaline Varied Varied — some euryhaline Euryhaline

slightly more southerly latitude during the Eocene (Smith et al., 1981), palaeogeographical reconstructions place the Hampshire Basin in a proto-English Channel (Curry, 1978; Gale et al., 1999). However, in the absence of constraints for wind climate, it is not possible to unequivocally estimate wave climate. Lateral variations in the middle Eocene sedimentary facies suggest that the Bartonian sea floor was wave-graded (Figs. 6 and 11; Plint, 1983), and intermittent shell and gravel lags throughout the succession at Alum Bay (Fig. 3) are consistent with some storm influence and moderate wave energy. Thus, assuming a moderate wave climate, the percentage of mud grade sediment is used to reconstruct the Bartonian palaeobathymetry at Alum Bay and Whitecliff Bay (Fig. 11). These estimates may be subject to the large uncertainties associated with estimating wave climate. For this reason, we plot possible end-member palaeodepth curves based on the low and high energy wave climate calibrations presented by Dunbar and Barrett (2005). However, it is not possible to place more precise errors on these estimates.

Magneto- Sedimentary chrons facies

)

h (m

flat discoidal planoconvex rounded planispiral ovate porcellaneous globular elongate trochospiral lenticular trochospiral subglobular trochospiral flattened elongate cylinderical biconvex keeled

G

60 G

Sandy–muddy clay and hard substrates Infaunal Epiphytic and sediment Muddy

A further complication of using the percent-mud palaeodepth proxy is the assumption that the b 63 μm sediment fraction contains only derived/transported clasts. Glaucony is present throughout the succession at Alum Bay, and is inferred to be in situ (Section 4.2). It was not possible to remove or estimate the percentage of the glaucony grains in the b 63 μm fraction. As a result, the percent-mud based palaeodepth estimates may be biased if there are significant variations in the in situ component. The colour reflectance of the bulk material (Fig. 7) is however, largely invariant throughout the succession, suggesting that variations in the glaucony component of the b 63 μm fraction are not significant. Water depth variations based on percent-mud in the succession at Alum Bay are consistent with the systems tracts inferred from foraminifera and sediment properties (Figs. 7 and 8). The amplitude of water depth variation based on percent-mud (~ 40 m, moderate wave climate) is smaller than estimates derived from the foraminiferal assemblages and morphogroups (~ 20–80 m). This discrepancy may

Foraminifera morphogroups

pt De

70

C18n

Substrate

Water depth

Salinity

Substrate

inner-central shelf

normal marine

central shelf

euryhaline

central shelf

euryhaline

inner-central shelf

normal marine

central shelf

euryhaline

muddy

normal marine

sandy-muddy

muddy

50

G

40

G G

30

G G

central-outer shelf G

C18r

20

outer shelf

G G

C20nr C19r

G

10

G G

0 0

10

20

30

40

50

60

70

80

Number of foraminifera per g sediment Fig. 10. Distribution of foraminiferal morphogroups in the Barton Clay Formation succession at Alum Bay. Variations in water depth, salinity and sediment substrate are based on the relationships summarised in Table 2.

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Hengistbury Head Sedimentary facies

Sedimentary facies

Relative water depth curve

% mud palaeodepth estimate (m) 100 80

+ 20 m

Whitecliff Bay

Alum Bay

-

60

40

20

Grain sorting 0

0

1

Grain sorting 2

0

1

2

% mud palaeodepth estimate (m) 100 80

60

40

20

Sedimentary facies 0

G

G G

Above fair weather wave base (~15m)

20 m

20 m

N. prestwichianus G

C18r

G G

G G

G G

G

10 m

10 m

10 m

G G

C19r

G C19n

G

C20n

G

0m

G G G 0m

0m

G

Base of Barton Clay Fm.

N. variolarius

Fig. 11. West–East transect of lower Barton Clay Formation successions within the Hampshire Basin. The successions are correlated on the basis of the basal conglomerate and the N. prestwichianus horizon. Sediment grain size is used to infer semi-quantitative water depth estimates for the successions at Alum Bay and Whitecliff Bay (both Isle of Wight) using the percent-mud proxy and the assumption of a moderate wave climate (symbols; low and moderate-high wave climates are used to provide end-member palaeodepth estimates and define the shaded area). Facies variations observed at outcrop level are used into infer water depth variations in the succession at Hengistbury Head. During the late Lutetian and early Bartonian (C20n–C18r), increases in water depth identified in the three successions of the Barton Clay Formation appear to have been synchronous.

reflect the limitation of the percent-mud proxy to depths shallower than the wave base (~ 70 m), and potential uncertainties estimating ancient wave climates. Water depth estimates for the succession at Whitecliff Bay are deeper than those for the successions at Alum Bay, and Hengistbury Head, and may indicate a location more distal from the palaeoshoreline. Trends in sediment grain size and sorting are mirrored in both lower Barton Clay Formation successions on the IOW, although water depth variations in the succession at Whitecliff Bay are slightly smaller. Both of the IOW successions are barren of foraminifera in their lower parts, limiting further comparison. During the late Lutetian and at least the early Bartonian, water depth variations inferred for the Barton Clay Formation at Alum Bay appear to be regional in extent (Fig. 11). In the neighbouring Belgium Basin, the equivalent Bartonian continental shelf succession, the Kallo Complex, contains evidence for three, third-order sedimentary cycles (Vandenberghe et al., 1998). These cycles may indicate significant provincial sea level variation during the middle Eocene. However, it is unclear how these cycles precisely relate to those in the Hampshire Basin because of the poor age constraints for the microfossil barren Kallo Complex (Vandenberghe et al., 1998).

6. Estimating eustatic variations Basin-wide correlation of the sedimentary facies in the lower Barton Clay Formation supports regional water depth variations of up to 40–60 m during the late Lutetian and early Bartonian. To investigate potential contributions from glacio-eustasy and facilitate the comparison with other records of sea-level change, estimates of water depth should be converted to eustasy.

The tectonic evolution of Southern England has been an active area of research for decades and the particulars of the exhumation history are still disputed (Chadwick and Evans, 2005; Daley and Edwards, 1971; Gale et al., 1999; Hester, 1965; Hillis et al., 2008; Jones, 1981; Plint, 1982). Subsidence in the Hampshire Basin commenced during the late Cretaceous under a compressional regime resulting from the progressive closure of the Tethys Ocean (Stonely, 1982). The switch from an extensional regime inverted the sense of motion on many of the regional scale normal faults, notably the Purbeck–IOW fault system extending from south Dorset to the eastern Solent (Fig. 1, Chadwick, 1986; cf. Insole et al., 1998 e.g. Stonely, 1982). The onset of the inversion on the Purbeck–IOW belt is debated, with estimates of Chalk denudation in the early Palaeocene ranging from 0 to 200 m (Gale et al., 1999; Hillis et al., 2008; Insole et al., 1998; Jones, 1981). As a result, two tectonic scenarios have been proposed for Palaeogene evolution of the Hampshire Basin. Based on presence of reworked Palaeogene and Cretaceous clasts and fossils in Eocene strata (Gale et al., 1999; Plint, 1982) and lateral variations in the thickness of early Oligocene sediments (upper Bembridge Limestone, Daley and Edwards, 1971), the first scenario invokes episodic uplift throughout the Palaeogene (Gale et al., 1999). The second scenario attributes the majority of the uplift to the climax of the Alpine Orogeny during the Miocene, citing the lack of thinning of Palaeocene through Oligocene strata over the crest of the Sandown monocline as evidence that compressional deformation and exhumation post-date the preserved strata (Chadwick and Evans, 2005; Hillis et al., 2008). Derived specimens of early Lutetian (NP14b) and Cretaceous calcareous nannofossils are observed in the lower Barton Clay Formation at Alum Bay (~ 18–25 m above the base). This evidence suggests that Eocene and Cretaceous material was exposed either within, or sourced from outside the basin during the early Bartonian.

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Given the proximity of Alum Bay and Whitecliff Bay to the Purbeck– IOW fault belt, and the presence of reworked clasts and fossils in these Bartonian successions (Aubry, 1983, 1986; Gale et al., 1999; Plint, 1982; this study), the evidence for Eocene tectonic quiescence in the Hampshire Basin is equivocal. The absence of the middle Barton Clay Formation in the succession at Whitecliff Bay (Gale et al., 1999) likely indicates that inversion on the Purbeck–IOW fault system was spatially and temporally variable, with the greatest uplift focused in the eastern part of the basin during the Bartonian. In the absence of a quantitative tectonic model for the Hampshire Basin, it is not possible at present to calculate eustatic variations for the Barton Clay Formation. Neither is it possible to unequivocally attribute the water depth variations inferred from the successions at Alum and Whitecliff Bay to either change in the volume of ocean water (i.e. glacioeustasy) or in the volume of the ocean basin. However, tentative correlations of the records from Alum Bay to sedimentary cycles on the East Tasman Plateau and NJCP support a eustatic origin for some component these cycles (Fig. 12).

some evidence to support warming at the high latitudes (Burgess et al., 2009; Houben et al., 2009), precise temperature and seawater δ18O contributions to the ‘MECO’ excursion have not yet been established at other locations (Bohaty and Zachos, 2003; Bohaty et al., 2009). Additionally, the ‘MECO’ perturbation is recognized in records of bulk carbonate δ13C and δ18O from the pelagic and hemipelagic Contessa and Alano successions respectively in Italy and has been linked with local shifts in terrigenous sediment and organic carbon supply (Jovane et al., 2007; Spofforth et al., 2010). A single-species δ18O record from the Barton Clay Formation at Alum Bay was generated using a gas source mass spectrometer at the University of Cambridge. Samples are based on a large number (15–60) of exceptionally well-preserved specimens of C. ungerianus. Rapid and clearly defined shifts in benthic δ18O of ~ 1‰ are observed at ~ 20 m (~ 40.45 Ma) and 57 m (~ 39.9 Ma) at Alum Bay (Fig. 13). The positive benthic δ18O excursion at ~ 20 m does not correspond to a notable change in water depth, occurring ~ 2 m above the ‘N. prestwichianus’ transgressive horizon. This δ18O excursion may reflect a local decrease in ambient water temperature and/or seawater δ18O related to the long-term (third-order) increase in water depth. In contrast, the negative δ18O shift at ~ 57 m coincides with a fourth-order sequence stratigraphic cycle, in which water depth is inferred to increase by ~ 20–50 m (Figs. 7 and 8; Table 1). This excursion is unlikely to reflect an ambient temperature change associated with the increase in water depth, but probably indicates a regional warming event and/or change in seawater δ18O. The seawater δ18O composition of shallow continental seaways is influenced by the mixing of riverine and marine waters, reflecting variations in the position of the palaeoshoreline. The 1‰ decrease in benthic δ18O at 57 m (~ 40 Ma) may reflect an increased contribution

7. Evidence of the ‘middle Eocene Climatic Optimum’ isotope excursion A large (0.8–1.5‰) negative excursion recently identified ca. 40.0 Ma in records of bulk carbonate and foraminiferal δ18O from Ocean Drilling Program (ODP) Sites in the Southern, Atlantic and Indian Oceans (Bohaty and Zachos, 2003; Bohaty et al., 2009), has been termed the middle Eocene Climatic Optimum (MECO). The ‘MECO’ isotope excursion is interpreted by some to reflect transient global warming of up to 6 °C, assuming ice-free conditions at this time (Bohaty and Zachos, 2003; Bohaty et al., 2009). Although there is

Hampshire Basin chronology

Sedimentary cycles: Alum Bay

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100 80 60 40 20

0

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0

Fig. 12. Correlation of water depth variations observed in the Barton Clay Formation to (1) the sequence stratigraphy of ODP Site 1172 (East Tasman Plateau, Rohl et al., 2004) and (2) the ‘best estimate’ sea level curve from New Jersey Coastal Plain (NJCP, Miller et al., 2005). Ages are quoted on the Berggren and Pearson (2005) GPTS. Fourth order sedimentary cycles at Site 1172 are based on our interpretation of the [Ca] record, using the same criteria as Rohl et al. (2004) to infer third-order cycles, i.e. minimum [Ca] as a result of reduced calcareous nannofossil content reflect sea level falls. Green dashed lines denote the correlation of systems tracts. Max flooding surfaces at ~ 40.6 and 39.9 Ma are observed in all the successions. Ages for Site 1172 and the NJCP are well constrained from magnetostratigraphy, typically better than 0.05 and 0.3 m respectively (Fuller and Touchard, 2004; Miller et al., 1990).

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from fresh riverine waters, but this is inconsistent with the inferred increase in water depth (Figs. 7 and 8) and presumed landward migration of the palaeoshoreline. The amplitude (1‰) and timing of the negative benthic δ18O excursion that commences at 57 m broadly coincides with the ‘MECO’ excursion identified at other localities (~ 40 Ma, C18n2n; Bohaty et al., 2009; Jovane et al., 2007; Spofforth et al., 2010). It appears that some component of the 1‰ benthic foraminiferal δ18O excursion observed in the Barton Clay Formation record reflects the same ‘MECO’ temperature/seawater δ18O perturbation that has been recorded open ocean sediments (Bohaty and Zachos, 2003; Bohaty et al., 2009). The contemporaneous increase in water depth in the Hampshire Basin, which we argue can also observed on the East Tasman Plateau (Fig. 12), indicates that some component of the ‘MECO’ δ18O excursion recorded in the Barton Clay Formation and possibly at other localities, could reflect a change in seawater δ18O related to a reduction in global ice volume. This is consistent with the physical and proxy data that support significant continental ice storage prior to this period (Browning et al., 1996; Dawber and Tripati, accepted; Edgar et al., 2007; Miller et al., 2005; Moran et al., 2006; St. John, 2008; Stickley et al., 2009; Tripati et al., 2005, 2008).

8. Summary Sequence stratigraphic evidence for moderate amplitude (40–60 m) relative sea-level variations in the Hampshire Basin during the late Lutetian and Bartonian better defines the global nature of middle Eocene sea level variations. If additional work can characterize and quantify the extent and spatial distribution of uplift within the basin during the

Bartonian, the well-preserved successions of the Barton Clay Formation at Alum Bay and Whitecliff Bay would be good candidates for the calculation of eustasy using backstripping. The youngest sequence stratigraphical cycle documented in the Barton Clay Formation at Alum Bay (ca. 40 Ma) coincides with a large amplitude negative excursion in the local benthic δ18O record. The timing and amplitude of this excursion is consistent with the ‘MECO’ δ18O event previously reported in open-ocean sediments (Bohaty et al., 2009). The new records suggest that in the Hampshire Basin, the MECO δ18O excursion may reflect a combination of warming and a reduction in seawater δ18O as a result of a transient decrease in global ice volume. Additional temperature and sea level reconstructions are needed to assess what the MECO δ18O excursion represents on a regional scale. Supplementary materials related to this article can be found online at doi:10.1016/j.palaeo.2010.12.012.

Acknowledgements The authors gratefully acknowledge M.P Aubry, L. Booth, S. Crowhurst, and M. Hall for providing the technical assistance, and I. West, J. Todd, D. Kemp and N. McCave for the discussions of this work. This manuscript benefited from the constructive reviews of editor F. Surlyk, G. Dunbar and an anonymous reviewer. We thank Natural England and the Needles Theme Park for access to the succession at Alum Bay. C.F.D. was funded by a NERC studentship. A.T. gratefully acknowledges the financial support by NERC, Magdalene College at the University of Cambridge, and the UCLA Division of Physical Sciences.

Barton Clay Formation, Alum Bay gra

M

Bio

hy

p gra

ati str

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n ag

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t ep

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ati str eto

)

m h(

D

y ph

-3.6

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-2.8

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0.6

ODP Site 738 Southern Ocean

18O

1.2

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0

-0.6

18O

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--40.0Ma

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40.5

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60

40.0

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70

BAR-2

--40.1Ma

0

C20n

N. prestwichianus

BAR-1

10

41.5 HST

N. rectus

--40.4Ma

NP16

20

C18r

30

MFS

42.0 TST

41.3Ma42.5Ma- --41.5Ma--

HST

42.5

MFS

Fig. 13. Comparison of the Barton Clay Formation sedimentary, stratigraphic and benthic foraminiferal δ18O records to published benthic foraminiferal oxygen isotope records from the Southern Ocean (Bohaty et al., 2009). The grey box denotes the correlation of the ‘MECO’ δ18O excursion.

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