2010 Paleo-Eo.pdf

167MAGNETOSTRATIGRAPHY SUSCEPTIBILITY CORRELATION

MAGNETOSTRATIGRAPHY SUSCEPTIBILITY USED FOR HIGH-RESOLUTIONCORRELATION AMONG PALEOCENE–EOCENE BOUNDARY SEQUENCES

IN EGYPT, SPAIN, AND THE U.S.A.

BROOKS B. ELLWOODDepartment of Geology and Geophysics, Louisiana State University,

E235 Howe-Russell-Kniffen Geoscience Complex, Baton Rouge, Louisiana 70803, U.S.A.e-mail: [email protected]

AZIZ M. KAFAFYGeology Department, Faculty of Science, Tanta University, Tanta 31527, Egypt

AHMED KASSABGeology Department, Faculty of Science, Assiut University, Assiut 71516, Egypt

JONATHAN H. TOMKINSchool of Earth, Society, and Environment, University of Illinois, 428 Natural History Building,

1301 W. Green Street, Urbana, Illinois 61801, U.S.A.ABDELAZIZ ABDELDAYEM

Geology Department, Faculty of Science, Tanta University, Tanta 31527, EgyptNAGEH OBAIDALLA

Geology Department, Faculty of Science, Assiut University, Assiut 71516, EgyptKELLI W. RANDALL

Marathon Oil Company, 5555 San Felipe Rd., Houston, Texas 77056, U.S.A.AND

DAVID E. THOMPSONOffice of Geology, MDEQ, P.O. Box 20307, Jackson, Mississippi 39289-1307, U.S.A.

ABSTRACT: The magnetostratigraphy susceptibility technique is used to establish high-resolution correlation among Paleocene–Eoceneboundary sequences in Egypt, Spain, and the U.S.A. This work initially focuses on the Global boundary Stratotype Section and Point (GSSP),defining the base of the Ypresian Stage (lowest Eocene), located in the Dababiya Quarry near Luxor in Upper Egypt. The base of the Eocenerepresents the beginning of the Paleocene–Eocene Thermal Maximum (PETM) identified by a negative carbon isotope (δ13C) excursion.While onset of the CIE is somewhat gradual in most reported Paleocene–Eocene (P–E) sections, at the GSSP it is very abrupt and beginsimmediately after an unusual lithologic change that magnetic susceptibility (MS) and other data indicate represents a short erosional ornondepositional hiatus. Comparison of MS zones from five well-studied marine sequences (the Dababiya Quarry GSSP, Jebal El Qreiya,also in Upper Egypt, Zumaia in northern Spain, Alamedilla in southern Spain, and the MGS-1 Harrell Core from southeastern Mississippi,U.S.A.) with that from the GSSP site shows a period of reduced sedimentation and nondeposition through the boundary interval in theGSSP. This interval, estimated to have lasted for ~ 10,000 years, is less than the biostratigraphic resolution for the site. Due to the hiatusin the GSSP, we have chosen the P–E section in Zumaia as the MS reference section for the P–E boundary interval. Because the correlationbetween the Zumaia section in Spain and the MGS-1 Core from the U.S.A. is excellent, and because the MGS-1 data set represents a longerinterval of time than does the Zumaia data set, we use the MS data from the MGS-1 Core to extend the MS zones from Zumaia and establisha MS composite reference section (MS CRS) for the P–E boundary interval sampled. Orbital-forcing frequencies for the Zumaia referencesection are then identified, via spectral analysis. Extending the MS zones into the MS CRS allows age assignment to MS zones for all fivesections with a resolution of ~ 26,000 years.

KEY WORDS: magnetostratigraphy susceptibility, Paleocene–Eocene, Global Boundary Stratotype, orbital frequencies

Application of Modern Stratigraphic Techniques: Theory and Case HistoriesSEPM Special Publication No. 94, Copyright © 2010SEPM (Society for Sedimentary Geology), ISBN 978-1-56576-199-5, p. 167–179.

INTRODUCTION

Global boundary stratotype sections and points (GSSPs) arecurrently being established to formally define the beginning ofeach geologic stage for the Phanerozoic (Salvador, 1994). Acompilation of the GSSPs ratified to date, as well as the support-ing data, can be found in The Geologic Time Scale 2004 (Gradsteinet al., 2004). However, the biostratigraphic and geochemical dataused to define the GSSPs are not always sufficient to make globalgeologic correlations possible. Most serious are the many prob-

lems in using fossils for high-resolution correlation, which in-clude facies differences, changing evolutionary rates, pseudo-extinctions, and the elimination of critical marker horizons byerosion or nondeposition. Other methods can be used to make upfor the limitations of biostratigraphy, but none are perfect. Isoto-pic dates are few and lack sufficient resolution to date GSSPs.Geochemical techniques are useful for correlations but are ad-versely affected by weathering, and some sequences, such asthose are below the carbonate compensation depth (CCD), oftenlack the carbonate needed for carbon and oxygen isotopic work.

BROOKS B. ELLWOOD ET AL.168

Magnetostratigraphic methods provide a useful supplemen-tary relative dating technique, and the Subcommission on Stratig-raphy includes formal usage of magnetic-property measure-ments in stratigraphy (Salvador, 1994). Magnetostratigraphypolarity, based on the remanent magnetization (RM) of rocks, isbeing routinely used for correlation in Cenozoic and Mesozoicrocks (e.g., Berggren et al., 1995; Gradstein et al., 2004), but it islimited to well-consolidated materials that can be oriented, usu-ally excluding cuttings and highly friable sediments. Anothermagnetostratigraphic technique exists, the magnetic susceptibil-ity (MS) method, and it can be used on a wide range of samples,including cuttings and friable outcrop materials. Thus it can beused in circumstances that limit the usefulness of other strati-graphic techniques, and in conjunction with biostratigraphy,isotope stratigraphy, and polarity stratigraphy it can providehigh-resolution relative dates. The method works even when therocks sampled are somewhat weathered and/or have been heatedto moderate temperatures. In some cases, for example measure-ments on Moroccan samples (Ellwood et al., 1999), conodontalteration indexes yield temperatures > 350° C). Furthermore, themethod is inexpensive, fast, and often unaffected by alteration ordepth effects, and requires only small samples that are relativelyeasy to collect, and fast and inexpensive to measure.

In this study, we use MS to characterize the Paleocene–Eocene(P–E) boundary. We chose this boundary because it represents asignificant extinction event associated with a major climaticwarming event, the Paleocene–Eocene Thermal Maximum(PETM). The PETM is now defined to begin with a carbon isotopeshift that is also used to define the beginning of the YpresianStage, the first stage in the Eocene Series. The Ypresian GSSP isone of the first to use carbon-isotope chemostratigraphy, insteadof a biostratigraphic marker, to define an important GSSP. Be-cause there may be sections where isotopic results are problem-atic, and the boundary is not resolved by polarity stratigraphy (itfalls within polarity chron C24r; Gradstein et al., 2004), weemploy the MS method to supply the high-resolution strati-graphic record needed for broad correlation. We demonstratehere how MS methods complement biostratigraphic data byusing the two together to produce a high-resolution floating-point time scale (FPTS) for the P–E boundary interval. This newtime scale indicates a previously unobserved ~ 10,000 year hiatusin sediment accumulation at the P–E boundary level in the UpperEgypt GSSP section.

SETTING

The Paleocene–Eocene Boundary Event

The P–E Series boundary marks the division between theThanetian and Ypresian stages and is dated at ~ 55.8 Ma (Gradsteinet al., 2004). It has been argued that during the earliest Ypresianthere was a massive, sudden input of ~ 1,200 to 2,000 gigatons ofcarbon as methane into the atmosphere, causing a significantglobal warming event (Norris and Röhl, 1999). This methanerelease could have been caused by the melting of frozen hydro-carbon accumulations (gas hydrates), primarily containing meth-ane on the ocean floor, or through other processes (Lourens et al.,2005), as a result of excessive volcanism (Storey et al., 2007), ordue to an astronomical triggering mechanism (Westerhold et al.,2007). The residence time of carbon in the atmosphere is ~ 120 kyr,which it has been argued resulted in long-term alteration ofglobal climate (Norris and Röhl, 1999). This change in the envi-ronment drove speciation and caused the extinctions, changesthat occur in the fossil record at and just above the P–E boundary.For example, the earliest horse species, Hyracotherium, evolved at

this time, but approximately 30–50% of all benthic foraminiferabecame extinct (Röhl et al., 2000). There was a nearly total reversalof ocean water current flow over a period of perhaps 100 kyr at theP–E boundary, from the Southern Ocean being the source of deepwater flow in the latest Paleocene to Northern Ocean sources inthe earliest Eocene (Nunes and Norris, 2006). The wide dissemi-nation of warm water to the deep sea allowed distribution oftropical climate into high latitudes. As a result there was a periodof at least 40 kyr when the carbon isotope ratio changed from pre-PETM levels to full excursion levels (a negative δ13C shift of 2‰or more), a change that supports the argument of release of gashydrate into the oceans and associated global warming (Röhl etal., 2000).

The GSSP for the Paleocene–Eocene(Thanetian–Ypresian) Boundary

The Dababiya Quarry section in Upper Egypt (E in Figure 1;Figure 2) was chosen as the GSSP and formally ratified in 2003(Gradstein et al., 2004). It is located in an inactive quarry, approxi-mately 1 km east of the village of Dababiya, on the east bank of theNile River, to the east of the Luxor–Aswan Road and ~ 35 kmsouth of Luxor. Sediments in the section represent an openmarine setting, possibly ranging from neritic to upper-bathyalenvironments (Ouda, 2003).

The Dababiya Quarry section was selected as the GSSP due tothe nearly complete biostratigraphic, chemostratigraphic, andlithologic records across the P–E boundary and is containedwithin the Esna shale, at the Esna 1–Esna 2 contact (Fig. 2). Withinthe Esna 1–Esna 2 P–E boundary interval is the Dababiya ChannelSequence (DCS, Fig. 2) composed of thin clay beds beneath aphosphatic coprolite-rich shale grading into thick fossiliferouscalcareous mudstone. The DCS grades into Esna 2, which is agenerally lithologically homogeneous clay, dark in color with nobedding. The DCS at Dababiya possesses a thin, basal dark grayclay marking the P–E boundary (Fig. 2).

The geochemistry of the Dababiya Quarry section shows anegative carbon isotope excursion (CIE) of -3‰ at the base of theDCS. Therefore, the base of this CIE has been chosen as the P–E

FIG. 1.—Paleogeographic reconstruction map (Blakey, 2008) forthe Thanetian–Ypresian (P–E) boundary time (~ 55 Ma) show-ing site locations. Samples used in this study are from UpperEgypt (E), Spain (A, Z), and the U.S.A. in Mississippi (M).They represent shallow marine areas in which sedimentswere accumulating during P–E time.


boundary marker event or Point. The CIE also correlates with amassive benthic foraminiferal extinction (Dupuis et al., 2003).Before 2003, the P–E boundary was traditionally defined basedon the highest occurrence of the planktonic foraminifer Morozovellavelascoensis (see Dupuis et al., 2003, for a detailed description ofthe lithostratigraphy, clay mineralogy, geochemistry, and pale-ontology).

The biostratigraphy of calcareous nannofossils and plank-tonic foraminifera indicates that the P–E boundary represents asignificant extinction event. Preservation of microfossils rangesfrom excellent to poor, depending on the lithology. Even so, themicrofossil record is sufficiently complete to show extinctionsbelow and the appearance of new species above the boundary.According to Dupuis et al. (2003), typical planktonic foramin-iferal assemblages for the top of Esna 1 within the DababiyaQuarry include Morozovella velascoensis, M. acuta, M. aequa, M.subbotinae, and Acarinina soldadoensis. The DCS is bracketed be-low by the last appearance (LA) of A. esnanensis, Igorinabroedermanni, A. angulosa, and A. wilcoxensis and above by the LAof Pseudohastigerina wilcoxensis (see Berggren and Ouda, 2003a,2003b; Ouda and Berggren, 2003). The calcareous nannofossilassemblage directly above the boundary includes the newly

evolved species Rhomboaster spp., Pontosphaera spp., Discoasteranartios, and D. araneus (see Dupuis et al., 2003). Many of thesemicrofossils are excursion taxa that develop and go to extinctionall within the CIE interval. The benthic foraminiferal assemblage,consisting of newly evolved species Anomalinoides aegyptiacusand A. zitteli, originating in the CIE interval’s benthic foramin-iferal extinction, suggests an outer-neritic to upper-bathyal ma-rine environment for the site (Dupuis et al., 2003).

The Paleocene–Eocene Boundary Section at Jebal El Qreiya

Information supporting interpretations from the GSSP comesfrom the Jebal El Qreiya section, located 110 km to the NNE of theGSSP. The geochemistry at Jebal El Qreiya supports evidence forglobal warming, with a negative CIE at the P–E boundary. As atthe GSSP, the CIE represents a negative 3‰ carbon isotope shift,within the range of normal CIE values for other Egyptian P–Esites (Knox et al., 2003). At the Jebal El Qreiya section, the shalewithin the P–E boundary layer contains the largest abundance ofkaolinite within the CIE, followed by a decline in kaolinite above(Soliman, 2003). Such increases in kaolinite indicate a warm, wetclimate during the CIE. There is also a higher proportion of

1.0E-07

Magnetic Susceptibility (m3/kg)

0.0

0.5

1.0

1.5

2.0

2.5

3.0

3.5

4.0

Paleocene–Eocene GSSPUpper Egypt

Yp4

Yp3

Yp2

Yp1

ThZ

ThY

Yp5

Yp6

Yp7

ThX

ThW

2.0E-073.0E-08

DCS

Ypresian

Thanetian

Thanetian

Ypresian

MSzone

LithLog

FIG. 2.—Outcrop photo and MS data from the P–E boundary GSSP in the Dababiya Quarry. (E in Figure 1). Raw MS data are dashed;smoothed data are represented by the solid curve; MS zones (with labels) are presented as bar logs, and methods are explainedin the text, where filled bars represent higher MS values and open bars are lower MS values. The location of the P–E boundaryis indicated. MS zones beginning with Yp and Th are Ypresian and Thanetian, respectively. Short black interval in the lith-logrepresents the Dababiya Channel Sequence (DCS; Dupuis et al., 2003) boundary layer extending across the outcrop (naturaldiscoloration in the figure) between the ESNA1 and ESNA2 shales.


chlorite and illite coincident with the CIE, which suggests anincreased supply of detrital clays.

The upper part of the CIE at Jebal El Qreiya, as at the DababiyaQuarry, is marked by changes in the abundance of planktonicforaminifera and the complete extinction of most benthic fora-minifera (Knox et al., 2003). However, the calcareous-nanno-plankton and benthic-foraminifera zonal boundaries are indis-tinct and disputed. This is due to poor preservation of calcareousnannoplankton in the Esna shale beds. A distinct succession ofassemblages of benthic foraminifera was recognized at Jebal ElQreiya: Angulogavelinella avnimelechi below the CIE, Anomalinoidesaegyptiacus in the lower portion of the CIE, and Bulimina callahaniin the upper portion of the CIE (Dupuis et al., 2003).

METHODS

Magnetic Susceptibility (MS): General Comments

All mineral grains are “susceptible” to becoming magnetizedin the presence of a magnetic field, and MS is an indicator of thestrength of this transient magnetism in a material sample. MS isvery different from remanent magnetism (RM), the intrinsicmagnetization that records the magnetic polarity of materials.MS in marine sediments is generally considered to be an indicatorof iron, ferromagnesian-mineral or clay-mineral concentration.In the very low inducing magnetic fields that are generallyapplied, MS is largely a function of the concentration and compo-sition of the magnetizable material in a sample. MS is a second-rank tensor (see Equation 1 and discussion below) and thereforehas an anisotropic signal. The anisotropy of magnetic susceptibil-ity (AMS) of most rock types has been well studied (e.g., Tarlingand Hrouda, 1993), and for most marine sedimentary samples isnot usually important. As a consequence, bulk (initial) measure-ments of MS are often performed without consideration of theAMS of samples. The anisotropic effect can be reduced whensamples are crushed before MS measurement. Our measure-ments are made on crushed samples.

Magnetizable materials in sediments include ferrimagneticminerals such as the iron oxide minerals magnetite and maghemite,and iron sulfide minerals, including pyrrhotite and greigite.These minerals acquire an RM (required for reversalmagnetostratigraphy). Other much less magnetic, paramagneticcompounds are present and may dominate the MS in marinesedimentary rocks. These include the clays, particularly chlorite,smectite, and illite, ferromagnesian silicates such as biotite, py-roxene, and amphiboles, iron sulfides including pyrite and mar-casite, and iron carbonates such as siderite and ankerite.

In addition to the ferrimagnetic and paramagnetic grains insediments, calcite and/or quartz is usually abundant, as are or-ganic compounds. These compounds exhibit a very weak negativeMS when placed in inducing magnetic fields, thus their acquiredMS is opposed to the low magnetic field that is applied. Thepresence of these diamagnetic minerals reduces the MS in a sample.Therefore, factors such as changes in biological productivity ororganic-carbon accumulation rates may be reflected in MS values.

Low-field MS, as used in most reported studies, is defined asthe ratio of the induced magnetic moment (Mi) or intensity ( Ji) ofmagnetization, to the strength of an applied, very low-intensitymagnetic field (Hj), where:

Ji = χijHj (mass-specific) (1)

or

Mi = kijHj (volume-specific) (2)

In these expressions, MS in SI units is parameterized as k,indicating that the measurement is relative to a one-cubic-meter volume (m3) and therefore is dimensionless; MS param-eterized as χ indicates measurement relative to a mass of onekilogram, and is given in units of m3/kg. Both k and χ exhibitanisotropy. Here we report mass specific MS without consid-eration of the AMS, and use χ to characterize this bulk (initial)low-field MS.

Presentation of MS Data

For presentation purposes and inter-data-set comparisons,the bar-log format, similar to that previously established forpresentations of magnetic-polarity data, is used. These bar logsare accompanied by both raw and smoothed MS data sets. Here,raw MS data (e.g., dashed line in Figure 2 illustrating data fromthe Paleocene–Eocene (Thanetian–Ypresian) GSSP in Egypt) aresmoothed using a smoothing spline (solid curve in Figure 2). Thefollowing bar-log plotting convention was then used. If the MSvalues during any cycle increase or decrease by a factor of two ormore, and if the change is represented by two or more data pointsin the splined data set, then this change is assumed to be signifi-cant and the highs and lows associated with these cycles aredifferentiated by filled (high MS values) or open (low MS values)bar logs (shown at the right in figures; e.g., Figure 2). This methodis best employed when high-resolution data sets are being ana-lyzed (large numbers of closely spaced samples) and helps re-solve variations associated with anomalous samples. Magnetic-susceptibility data are represented in log plots because of therange of MS values often encountered.

MS Measurements

All measurements reported here were performed using thesusceptibility bridge at Louisiana State University (LSU). Thisbridge was built by Marshall Williams and is a one-of-a-kindinstrument. It is calibrated using standard salts for which valuesare reported in the Handbook of Chemistry and Physics (Weast,1982) and by Swartzendruber (1992). We report MS in terms ofsample mass because it is much easier and faster to measure withhigh precision than is volume (see Ellwood et al., 1988). Eachsample is measured three times, and the mean and standarddeviation of these measurements is calculated. The mean of thesemeasurements is reported here. We also performed a number ofthermomagnetic-susceptibility measurements on samples mea-sured for MS, using the AGICO KLY-3S instrument, and thesedata are reported and discussed below.

PREVIOUS WORK ONMAGNETOSTRATIGRAPHY SUSCEPTIBILITY

Studies of Marine Sedimentary Sequences

Marine sediments are composed mainly of terrigenous, de-trital, components, and a biogenic component of the carbonateand/or siliceous tests of marine organisms. There is also acomponent of sediment derived from marine weathering pro-cesses, but this is generally minor. The MS signature mainlyrepresents contributions from the terrigenous and biogeniccomponents. Because the biogenic component is diamagnetic,in general, the terrigenous component dominates the MS. There-fore, the processes controlling the influx into the marine envi-ronment of terrigenous components usually account for the MSvariations observed (Ellwood et al., 2000; Ellwood et al., 2008).Thus, climate cycles that cause erosion and transport of terrig-


enous material result in a cyclic signal in the MS recorded in thedeposited sediments. There are many factors that can modifythis signature, but while the expected variations may be re-duced in magnitude, usually the character of the cycles remainsand can be extracted. Superimposed on these cyclic variationsare contributions from unique events, such as volcanic erup-tions or meteorite impacts (Ellwood et al., 2003), which may beof local, regional, or global significance.

Previous Results from LithifiedMarine Sedimentary Sequences

During diagenesis the MS of marine sediments decreases.This occurs because ferrimagnetic grains that dominate in un-lithified sediments are generally converted to paramagnetic phasesduring burial and diagenesis, mainly by the action of bacterialorganisms. This process transforms much of the iron to paramag-netic phases such as iron sulfide, carbonate minerals, and clays.In general, however, this process does not remove the iron fromthe system, so that total iron content is conserved. These newphases then contribute, along with the other paramagnetic detri-tal components present such as clays, and the remaining authi-genic ferrimagnetic constituents, to the MS in lithified sediments.MS values for most marine limestone, marlstone, shale, siltstone,and sandstone samples range from ~ 1 x 10-9 to ~ 1 x 10-7 m3/kg.Shale and some marlstone has the greatest AMS, but for most ofthese samples it is less than 2%. For limestone, marlstone, silt-stone, and sandstone samples the anisotropy is usually less than1%. Samples of shale and marlstone are friable, and measurementis generally on small broken fragments, further reducing theanisotropy effect. The error introduced by the AMS is thereforeminimal.

A study of the MS variability in a single limestone bed hasdemonstrated MS consistency over long distances (~ 25 km; workby Ellwood et al., 1999). Other work has demonstrated correla-tions over much greater distances (Ellwood et al., 2006; Ellwoodet al., 2008). Diagenesis also modifies some of the paramagneticconstituents that make up the terrigenous component of marinesediments. In many cases, clay minerals dominate much of thedetrital component, and these are generally not destroyed butmay be altered. The iron is often conserved in the secondary claysand pyrite, as is the paramagnetic behavior of these materials.Other common detrital components observed in marine rocks arethe ferromagnesian minerals, including biotite, tourmaline, andsimilar minerals (see Ellwood et al., 2000, for a discussion).Ultimately, following diagenesis, the MS observed for most ma-rine sedimentary rocks can be used as a proxy for physicalprocesses, primarily climate, responsible for delivering the detri-tal component of these sediments into sedimentary basins (e.g.,van Dobeneck and Schmieder, 1999; Weedon et al., 1999;Westerhold et al., 2008; Ellwood et al., 2008).

Cycles of Sea Level and Climate and other MS variations

Results on magnetic susceptibility for almost all studies ofmarine sedimentary rocks show many levels of cyclicity. Cer-tain cycles are interpreted to result from climatically drivenprocesses (van Dobeneck and Schmieder, 1999; Weedon et al.,1999; Westerhold et al., 2001; Ellwood et al., 2008). Other cyclesare very long-term and are interpreted as resulting from trans-gressive–regressive (T–R) trends due to sea-level rise and fallassociated with eustasy. As a general rule, during transgressivecycles, especially in distal marine sequences, MS is observed todecrease. This happens because detrital sediments are usuallytrapped near shore during transgressions. However, there are

distinct MS peaks that are correlated to maximum-floodingsurfaces (MFS), as well as to sudden influxes of detrital materialinto sedimentary basins by turbidity currents or other sedimentsuspensions. Sequence stratigraphic studies have demonstratedthat at times when sea level is at a maximum at the end of atransgressive cycle, producing a MFS, there is commonly amarked reduction in carbonate sedimentation, accompanied byan influx of detrital material into the marine environment. Thereis a corresponding MS high associated with these MFSs. Duringregressions, when base level is lowered due to falling sea level,erosion flushes detrital sediment into ocean basins and in-creases the MS of basinal sediments. Thus, T–R cycles play animportant role in creating some observed cyclic MS variations.The term MSEC was coined by Crick et al. (1997) to describe thephysical processes that produce MS trends. Results of this workhave shown that the MS is mineral dependent but may beindependent of the macro-lithology. Reworking of sedimentsby ocean currents and the variable deposition of clays mayresult in large MS changes in either limestone or shale, eventhough the lithology may appear to be relatively invariable. Inaddition, while local subsidence may cause a lithologic changefrom shale to limestone or vice versa, the MS often remainsrelatively constant. Because of their excellent biostratigraphiccontrol that includes related sections, many of our MS studieshave concentrated on GSSPs and related sections that have beenagreed upon by the International Geological Congress’s Sub-commission on Stratigraphy. The P–E boundary provides agood example of this effort.

SAMPLING AND RESULTS

Here we report results for Paleocene–Eocene boundarysamples from two localities in Upper Egypt, two localities inSpain, and one locality in the U.S.A. Descriptions of these locali-ties and the MS data acquired from them are presented below. Inall sections the MS data were smoothed using smoothing splinesand MS zones were developed using the rules discussed above.

Paleocene–Eocene (Thanetian–Ypresian) GSSP in theDababiya Quarry, Upper Nile Valley, Egypt

The GSSP section at Dababiya was sampled at intervals rang-ing from 20 to 50 mm, bracketing the boundary (Fig. 2). A total of90 samples were taken from 3.8 m of section. Magnetic-suscepti-bility results are presented in Figure 2. They exhibit low valuesand relatively low-frequency cyclicity in the latest Paleocene,with a sharp increase in value observed for the first sample in theEocene (Ypresian Stage), immediately above the base of the DCS.This sharp MS increase is coincident with the onset of the CIE inthe section over a 60 mm interval (three samples) and is followedby a rapid decrease. There is then an MS pulse with cyclicity thatreaches the highest MS values in the section at a point ~ 0.6 mabove the boundary. MS cycles then decay in amplitude, finallyreaching values that are broadly similar to those in the latestPaleocene.

The smoothed MS data were used to construct a sequence ofMS zones for the section, and these are shown in Figure 2,together with the position of the P–E boundary. MS zones arelabeled up-section beginning with the P–E as Ypresian Yp1 toYp7, the top of the interval we collected. MS zones below theboundary end in the uppermost Thanetian with ThZ, decreasingto ThW. Eventually, when the MS zonation for the entire ThanetianStage is established, MS zones will be numbered. This system ofnumbering for all sections reported here (Figs. 2 to 6) is tied to theGSSP nomenclature.


Paleocene–Eocene Section,Jebal El Qreiya, Upper Nile Valley, Egypt

The Jebal El Qreiya section of the central Nile Valley is locatedat the southern end of Jebal Abu Had, ~ 50 km northeast of thetown of Qena and ~ 110 km NNE of the GSSP at Dababiya. Thesection is east of the Nile River and north of the Qena–SafagaRoad in the Eastern Desert, at Latitude 26° 21' N and Longitude33° 01' E. As at the GSSP, the beds dip gently eastward and thesection contains the same lithologies represented in the DababiyaQuarry. The Jebal El Qreiya P–E boundary was sampled at 50 mmintervals, and the MS was measured on 72 samples from ~ 3.6 mof section. MS shows cyclicity throughout (Fig. 3). In the upper-most Paleocene, MS values exhibit a pronounced decrease, abovewhich there is a sharp increase to the highest MS values observedin this section (Fig. 3), but the magnitude is less than that exhib-ited in the GSSP section (Fig. 2). In the lowermost Eocene, MScontinues to show cyclicity with a slight overall upward decrease.MS values in the Paleocene (Thanetian) exhibit essentially thesame magnitudes in both the Jebal El Qreiya and Dababiyasections, but the large MS event immediately above the boundary

in the GSSP section (Fig. 2) lies slightly higher in the Jebal ElQreiya section (Fig. 3).

Paleocene–Eocene Section, Zumaia, Spain

The town of Zumaia is located on the southern coast of theGulf of Vizcaya, in north-central Spain (Z in Figure 1), and isconsidered ideal for MS study because it has been shown to be“one of the most expanded and complete marine sectionsacross the P–E boundary” (Schmitz et al., 2004). The sectionwas sampled at 50 mm intervals, and the MS of samples wasmeasured at LSU. Outcrop and MS data for the section areshown in Figure 4. Lithologies and δ13C values for the sectionare reported by Schmitz et al. (2004). We have subdivided theupper Thanetian limestone interval of Schmitz et al. (2004) intofive distinct beds (Fig. 4). Fairly regular MS cyclicity is ob-served throughout the section. In the upper part of the Pale-ocene there is a significant decrease in MS beginning in themarls below Bed 1. This decrease flattens out in the top oflimestone Bed 1 and shows little variability into the middle ofBed 3. Here, values begin to rise toward the P–E boundary,located at the top of Bed 5 (Fig. 4). In the lower Ypresian, the MSexhibits cyclicity around values observed in Bed 5, and thenvalues fall at the top of the section.

Schmitz et al. (2004) place the PE boundary at the top of Bed5, where the carbon isotopes begin the negative shift into theCIE. They then place the benthic-foraminifera extinction eventsignificantly above the top of Bed 5, well into the CIE (Schmitzet al., 2004). They also measured δ13C for the Dababiya Quarrysection, and they and others show that the benthic-foramin-ifera extinction event and negative δ13C shift in the DababiyaQuarry begin at the same level (Dupuis et al., 2003; Schmitz etal., 2004).

Paleocene–Eocene Section, Alamedilla, Spain

The town of Alamedilla is located about 60 km NNE ofGranada in southern Spain (A in Figure 1). The section wassampled by Birger Schmitz at 0.28 m average intervals (103samples) and sent to us for measurement. Samples close to theP–E boundary interval were collected at ~ 0.1 m spacing. TheMS data are somewhat scattered in the 15–18 m interval in thesection (Fig. 5), and thermomagnetic measurement indicatesthat some samples are weathered and contain ferrimagneticpedogenic components that have slightly affected the MSresults (discussed below). However, the smoothed MS trendsare consistent with those at the other P–E sections measuredand show a major MS peak at the P–E boundary, followed bya second significant peak ~ 3.5 m above the boundary (Fig. 5).One problem with the results from the Alamedilla section isthe coarse character of the MS zonation below the P–E bound-ary versus that above the boundary (Fig. 5). We believe thatthis reflects, in part, significant weathering in samples col-lected at this locality.

The Alamedilla section is clearly much expanded relativeto the Zumaia and Egyptian P–E boundary sections that wehave examined. Lithologies and δ13C values for Alamedillasection are reported by Lu et al. (1996; 1998). δ13C data show asteady, negative value for much of the section, and as inZumaia, the CIE begins before the benthic foraminifera extinc-tion event. Here, δ13C values associated with the CIE show arange of 2‰ (Lu et al. (1996) and Lu et al. (1998). Schmitz(personal communication) estimates the boundary to lie at thebenthic foraminifera extinction event that here occurs abovethe base of the CIE.

FIG. 3.—MS data from Jebal El Qreiya P–E section located in Egyptat (E in Figure 1). Symbols and notations are as in Figure 2. Thelocation of the P–E boundary is represented in the diagram intwo ways. Because of biostratigraphic uncertainties, litholo-gies similar to those at the GSSP are used to first assign the P–E boundary (solid line). Then the MS high value (dashed line)is used to represent a second possible choice for the locationof the P–E boundary (see text for discussion).


Paleocene–Eocene MGS-1 Core,Southeastern Mississippi, U.S.A.

The MGS-1 Harrell Core was drilled by the Mississippi Geo-logical Survey in southeastern Mississippi, USA (M in Figure 1)and is currently archived by the Survey. This core was sampledat ~ 50 mm spacing, where 235 samples were removed over a 5 minterval. Lithologies, pollen biostratigraphy, and δ13C values arereported by Harrington et al. (2004). The MS results show rela-tively low-amplitude cyclicity throughout most of the core, withthe exception of two major peaks, one at ~ 150.8 m depth and theother at ~ 149.1 m depth (Fig. 6). Lithologies, pollen biostratig-raphy, analysis of kaolinite clay, magnetostratigraphy polarity,and δ13C give an approximate position for the P–E boundary(Harrington et al., 2004) within the interval that we sampled. Theconfiguration of this data set is very similar to that observed forsamples from Alamedilla, Spain (Fig. 5). Based on this correlation,we have placed the P–E boundary at ~ 150.8 m depth in the MGS-1 Core, as indicated by the dashed line in Figure 6. This boundarypick is compared to the Zumaia MS results (Fig. 4) below.

Thermomagnetic Measurements

Thermomagnetic susceptibility was measured on a number ofsamples from the Zumaia P–E GSSP, MGS-1 Core, USA, and the

Alamedilla sections (a to d, respectively, in Figure 7). This in-volves heating the sample from room temperature to 700° C whilemeasuring the MS as the temperature rises. Paramagnetic miner-als show a parabola-shaped MS decay during this process be-cause the MS in these samples is inversely proportional to tem-perature of measurement (e.g., Hrouda, 1994). Ferrimagneticminerals, on the other hand, usually show an increase in MS upto a point where the MS decays toward the Curie temperature ofthe minerals responsible for the MS. During measurement, someiron-containing paramagnetic minerals such as illite (Ellwood etal., 2007) are unstable at higher temperatures and exhibit chemi-cal changes, often producing diagnostic peaks that result from theformation of a ferrimagnetic phase during this change. This is animportant result because these minerals usually denote a detritalor early diagenetic component (e.g., siderite) that was derivedfrom alteration of ferrimagnetic phases during diagenesis.

Thermomagnetic measurements for most samples in thisstudy show the parabolic decay (Fig. 7A–C) during heating,typical of the presence of paramagnetic grains (e.g., Hrouda,1994). However, in the Alamedilla samples (Fig. 7D) there is adistinct ferrimagnetic component that dominates most of thosesamples. In most of the Alamedilla samples there is both amagnetite (~ 580° C) and a maghemite (~ 610° C) Curie tempera-ture. We interpret this anomalous result to indicate that theseferrimagnetic constituents are the result of weathering in the

FIG. 4.—Outcrop photo and MS data from the Zumaia P–E section (Z in Figure 1). Symbols and notations are as in Figure 2. Bednumbers represent individual limestone beds (photo) located immediately below the boundary.


Alamedilla samples. In most of the samples measured fromZumaia (Fig. 7A) and Mississippi (Fig. 7C) there is a conversion ofparamagnetic phases, probably clays, at higher temperatures toferrimagnetic iron phases. This occurs in the laboratory becausethese paramagnetic phases, which are unstable at higher tempera-tures, have not previously been altered or destroyed by weatheringor other geological processes. GSSP samples have the lowest initialMS values, and both paramagnetic and ferrimagnetic componentsare present (Fig. 7B). Due to the very proximal setting for the MGS-1 Core, the magnitude of thermomagnetic susceptibility in thesesamples is higher than at the other sites reported here.

DISCUSSION

Correlation Among Sections

In most studies of marine sedimentary rocks where geologicalcorrelation is attempted, there are many uncertainties. Commonly,the most important data set used is biostratigraphic, but results canbe ambiguous because of problems due to facies-dependent organ-isms, poor recovery, unrecognized unconformities, and Lazarustaxa. This makes it important to employ independent correlationmethods in an effort to resolve these ambiguities. MSEC is one suchmethod. It requires fair to good biostratigraphic control to initiallydevelop a chronostratigraphic framework within which distinc-

tive MS zones can be directly correlated among sections with highprecision, even when biostratigraphic uncertainties or short un-conformities are known to exist.

We address the problem of correlating among sections byusing graphic comparisons between MS zones for sections re-ported here (Figs. 2–5) for which the biostratigraphy has beenwell determined. For these comparisons we selected the Zumaiasection as the reference standard, because (1) it is a well studiedsection for which biostratigraphic and carbon isotope data areavailable, and (2) it has been shown to be a complete and ex-panded marine section across the P–E boundary (Schmitz et al.,2004). In contrast, the GSSP and other Egyptian sections eachsuffer from problems, including a possible hiatus at the boundaryin the Egyptian sections and possible alteration.

First, we compared the Zumaia reference MS zonation toboth the GSSP and Jebal El Qreiya sections in Egypt (Fig. 8). Theresults for the boundary interval yield four line-of-correlation(LOC) segments with reasonable correlation between sections,but with significant changes in sediment accumulation rates asthe P–E boundary is approached. LOC2 represents a drop insediment accumulation rate, especially in the GSSP, which isthen followed by nondeposition (dashed line in Figure 8). Sedi-ment accumulation rates begin to increase above the boundary(LOC 3 in Figure 8) and then reach levels in LOC4 similar tothose in LOC1.

FIG. 6.—MS data from the MGS-1 Harrell Core (M in Figure 1).Symbols and notations are as in Figure 2. The dashed line isour pick for the location of the P–E boundary in the core basedon foraminifera data that define the boundary interval, andour comparison of the MS data with those data from theAlamedilla section (Fig. 5).

FIG. 5.—MS data from the Alamedilla P–E section (A in Figure 1).Symbols and notations are as in Figure 2. Samples wereprovided by Birger Schmidt (Schmitz et al., 2004; Schmidt,personal communication). The position of the P–E boundaryis shown.


A possible hiatus at the P–E boundary at Dababiya is sug-gested by the fact that the benthic-foraminifera extinction eventis located at the boundary and therefore at the base of the CIE inthe GSSP section. This is not the case in Zumaia, where thebenthic-foraminifera extinction event is located above the bound-ary and above the onset of the CIE. The MS character for theboundary interval also shows an abrupt offset at the boundary inthe GSSP (Fig. 2) that is typical for a hiatus.

High-resolution data sets from closely spaced samples canprovide means to identify unrecognized unconformities or ex-treme, rapid variations in depositional environments in marinesedimentary sequences. These hiatuses are reflected in geologi-cally instantaneous shifts in the stratigraphic distribution of thephysical or chemical properties measured. When nondepositionor erosion occurs, elements of cycles are not recorded or areremoved. This is then reflected in an offset in MS that does notrecord the expected cyclic variation that otherwise would havebeen observed in the data set. This is illustrated in the modelpresented in Figure 9. One such offset is observed in the P–E GSSPdata set, precisely at the boundary level (Fig. 2). The rapid change

in MS between the last Thanetian and the first Ypresian sampleseither represents a geologically nearly instantaneous physicalevent that has significantly increased detrital elements beingdeposited just above the boundary, or represents a sedimentaryhiatus or erosion.

An abrupt shift at the P–E boundary could be attributed to abolide impact (Cramer and Kent, 2005), but no physical evidenceknown to result from such impacts has been found (Schmitz et al.,2004). Work on the P–E boundary from other locations, includingDSDP and ODP sites (Norris and Röhl, 1999; Zachos et al., 2001;Tripati and Elderfield, 2005; Thomas and Bralower, 2005; Nunesand Norris, 2006) and marine and nonmarine outcrop sections(Röhl et al., 2000; Wing et al., 2005) demonstrates that the shifts inδ13C and other geochemical properties, including elemental ironand calcium concentrations, may have been more gradual andnot significantly pronounced as is the δ13C shift in the DababiyaQuarry GSSP section. Compounding this problem is the fact thatdue to rapid shoaling of the calcite compensation depth (CCD) atthe boundary, ODP sites from below the CCD did not depositcalcite, and the boundary interval is represented by clay (Zachoset al., 2005). Therefore, at such sites the first samples where calciteis available for carbon isotope analysis are above the boundary.Thus δ13C results at these sites generally show a rapid offset inδ13C that appears to produce the same pattern as that seen in the

FIG. 8.—Graphic comparison of the P–E GSSP section (Fig. 2) andthe Jebal El Qreiya section (Fig. 3) to the Zumaia (P–E)section (Fig. 4). Axes are labeled with the MS zone logs andnotation. Filled circles represent correlation points of corre-sponding MS zone boundaries between the GSSP (tied to theJebal El Qreiya data set by tie lies) and the Zumaia sectionthat we have chosen as the reference section for the P–Eboundary interval. Lines of correlation (LOC) are fitted tothe data. Dashed line represents the inferred hiatus at the P–E boundary GSSP and Jebal El Qreiya sections. Dashed lineswith arrows represent projections from the reference sectiontoward the GSSP data set.

FIG. 7.—Plots of thermomagnetic susceptibility for samples fromthe A, C) Spanish, B) Egyptian, and D) U.S. localities. Thesedata are typical for marine shale, marl, limestone, and silt-stone samples and indicate a significant paramagnetic com-ponent responsible for most of the low-field MS reportedhere. The exception is the Alamedilla, Spain, section (D)where the ferrimagnetic component dominates the MS. Thepeaks observed at higher temperatures, > 400° C, are duemainly to the breakdown of paramagnetic carbonates andclays to a ferrimagnetic phase (magnetite and maghemite).


GSSP section. This result is due to the hiatus in calcite depositiondue to CCD shallowing and produces a δ13C pattern like thatillustrated in Figure 9B.

Because the carbon isotope and biostratigraphic data aresomewhat ambiguous in the Jebal El Qreiya section, it is possiblethat there is no hiatus at the P–E boundary in that section (Fig. 3).Given the results from Zumaia, this would require that theboundary level is actually within MS zone Yp1. To evaluate this,we have adjusted the P–E boundary to lie at the level of thedashed line in Figure 3, and using the MS zones from the section,in conjunction with the new placement of the P–E boundary level,we have compared Jebal El Qreiya to the Zumaia referencesection (Fig, 10). The fit is better to Zumaia, requiring only threeLOC segments. This comparison indicates that while there is stillreduced sediment accumulation across the P–E boundary at JebalEl Qreiya, there may be no hiatus in sedimentation.

Comparing the Alamedilla and Zumaia sections yields anunusual set of LOC segments (Fig. 11). Given the distribution ofMS zones for the latest Paleocene at the site, there appears to bea radical increase in sediment accumulation rate in the Alamedillasection relative to the Zumaia section. This is followed by signifi-cantly reduced sedimentation through the boundary (LOC2; orsignificantly expanded section at Zumaia) and then a steppedrecovery above the boundary (LOCs 3 and 4, Fig. 11). We believethat, at least in part, weathering in the Alamedilla section (dis-

cussed above) has blurred the MS zonation at the site, makingassignment of MS zones problematic for the section. However,given the higher-frequency MS characteristics for the boundaryinterval and the general consistency of the splined data for theEocene data set (Fig. 5), we believe that the two MS peaks areprobably real, although assignment of MS zones again is prob-lematic. In addition, the magnitudes of these two peaks aresimilar to those of the two peaks we observe in the MGS-1 data set,suggesting that the lowest peak in that core may represent the P–E boundary. To test this, we graphically compared the MGS-1 MSzonation with the Zumaia reference section; those results arepresented in Figure 12. The result is an excellent fit for the MSzonation developed for two sections. This fit requires only twoLOC segments both of which show similar sediment accumula-tion rates, with a slight increase in rate for the Mississippi site,relative to Spain, in the lowermost Eocene. Due to this excellentfit, we decided to use the MGS-1 Core MS zonation to extend theZumaia reference section, and we projected six new MS zoneboundaries from MGS-1, through the two LOC segments and intothe reference section (dashed lines in Figure 12), thus creating aninitial MS composite reference section (MS CRS).

Times-Series AnalysisUsing the Fourier Transform (FT) Method

Fourier methods were used to determine the characteristicfrequencies of the Zumaia reference MS raw data set (unsmootheddata from Figure 4). We first assumed that the distance between

FIG. 10.—Graphic comparison of the Jebal El Qreiya sectionusing the MS high value as the boundary pick (dashed linein Figure 3) to the Zumaia (P–E) MS zone variation (Fig. 4).Axes are labeled with MS zone logs and notation. Filledcircles represent correlation points of corresponding MSzone boundaries between the Jebal El Qreiya data set and theZumaia reference section. Lines of correlation (LOC) arefitted to the data.

FIG. 9.—Hypothetical diagram illustrating A) a smoothly chang-ing data set (height and MS value) versus B) the same data setwhere erosion or nondeposition has removed 10 height units.This produces a hiatus (dashed line) in (Part B).


samples is linearly related with time, i.e., Dx is proportional to Dt,so that the FT method could be used. The more this assumptionis violated, the greater the amount of noise that is produced in thespectral graph. The data were then leveled and detrended, aprocess that reduces distortion in the final plot due to drift in thedata (Jenkins and Watts, 1968). The Fourier plot (Fig. 13) exhibitstwo well-defined spectral peaks and a third area of above-aver-age spectral power.

Our spectral analysis started by assigning Milankovitcheccentricity (E2, ~ 100 kyr) to the peak observed at 1 cycles/m inFigure 13, then the obliquity (O2, ~ 52.0 kyr and O1, ~ 39.8 kyr)and precession (P1, ~ 18.7 kyr) cyclicities (extrapolated fromBerger et al., 1992, for the Paleogene at ~ 55 Ma) coincide withregions of higher spectral power at ~ 2.3, ~ 2.8, and ~ 6.0 cycles/m, respectively (shaded bands in Figure 13). This compares wellwith previous work on the P–E boundary climate cyclicity. Forexample, Westerhold et al. (2007) showed an ~ 19 kyr precessioncomponent in their data from South Atlantic ODP cores, andwell defined ~ 400 and 100 kyr eccentricity peaks. In addition,Röhl et al. (2000) demonstrated that the P–E boundary intervalexhibits an ~ 21 kyr year cyclicity in elemental Fe and Ca forODP Site 690.

The MS bar logs calculated from smoothing splines yield acyclicity of 2.1–2.3 cycles/m (arrow in Figure 13), roughly consis-tent with an O2 obliquity signature for Zumaia. Assuming thatthese cycles, also observed for the raw MS data (Fig. 13), represent~ 26 kyr cyclicity at Zumaia, then the MS CRS represents ~ 500 kyrof time (Fig. 12). Given that the MS zonation at Zumaia iscorrelated among all the sites represented in this study, and giventhat this cyclicity is assigned an ~ 52 kyr/cycle age, we can assignrelative time estimates for MS zone boundaries and establish afloating-point time scale (FPTS) for the P–E interval representedby the MS CRS (Fig. 12). Therefore, each MS zone represents a halfcycle, or ~ 26 kyr (Fig. 12). The MS zones observed are variable inlength because sediment accumulation rates change naturally inthese sections, but the MSEC method is sufficiently robust to

accommodate relatively slight variations in accumulation rate orshort hiatuses.

Once the FPTS has been established, it is possible to assignages to the MS CRS zonation. Gradstein et al. (2004) report an agefor the P–E boundary of 55.8 ± 0.2 Ma. Because the total MS CRSzonation reported in Figure 12 represents ~ 500 kyr (each MS zonerepresents ~ 26 kyr), the base of MS zone ThQ (lowest defined MSzone) began at ~ 56.06 Ma. The entire MS CRS lasted until ~ 55.54Ma, the beginning of MS zone Yp10. Returning to the GSSP inUpper Egypt, we see that the lower half of MS zone Yp1 is missing(Fig. 8), representing a hiatus of at least 10,000 years. In addition,MS zones ThZ, the top of Yp1, and at least Yp2 have a significantreduction in sediment accumulation rate in the GSSP (Fig. 8).From this we conclude that there is from 10,000 to perhaps 20,000years of sedimentation missing (or it is very condensed) from theboundary interval at the Dababiya Quarry GSSP. Much of thisinterval may be represented by the thin DCS boundary layershown in Figure 2.

CONCLUSIONS

We used the MSEC method (Crick et al., 1997) to characterizefive Paleocene–Eocene (P–E) boundary sections, from Egypt(including the GSSP), Spain, and the United States. We find thata MS zonation established for these sequences can be correlatedamong all sections. Using the well studied section at Zumaia,Spain, as a reference standard, and correlating this reference

FIG. 11.—Graphic correlation between the Alamedilla (Fig. 5) andZumaia reference section (Fig. 4). Symbols are as in Figure 10. FIG. 12.—Graphic comparison between the MGS-1 core (Fig. 6)

and the Zumaia reference section (Fig. 4). Symbols are as inFigure 7. Solid lines in the diagram give the P–E boundary.LOCs fitted to the data are extended, and MS zone boundariesnot recognized in the reference section are projected throughthe LOC extensions and into the reference section to create anMS composite reference section (MS CRS). LOC 1 fitted to theThanetian data; LOC 2 fitted to the Ypresian data. A floating-point time scale (kyr) zonation is given. Ages are from cyclesderived in Figure 13.


with the longer P–E boundary section in Mississippi, we estab-lished an MS composite reference section (MS CRS) that extendsthe Zumaia MS zonation both above and below the P–E bound-ary. Using the Fourier Transform (FT) method, we have identi-fied Milankovitch periods present in the raw MS data, and theseallow us to establish a floating-point time scale (FPTS) for theMS CRS. Our results show that the MS CRS represents ~ 500,000years, extending from ~ 56 to ~ 55.5 Ma. The FPTS then allowshigh-resolution estimates of events recorded in the sedimentsdeposited during this interval of time. One result is that thereappears to be a hiatus in sediment accumulation at the P–Eboundary in the GSSP section in Upper Egypt, with a durationof from 10, 000 to 20,000 years.

ACKNOWLEDGMENTS

This project was funded in part by a U.S. National ScienceFoundation–Egypt Joint Science and Technology Board grant toBBE and AMK, grant #OTH6-008-002. We are grateful to BirgerSchmidt for providing samples from Alamedilla, Spain. SueEllwood is gratefully acknowledged for her help in sampling allsections. Thanks are extended to the reviewers for their construc-tive critical comments.

REFERENCES

BERGER, A., LOUTRE, M.F., AND LASKER, J., 1992, Stability of the astronomicalfrequencies over the Earth’s history for paleoclimate studies: Science,v. 255, p. 560–566.

BERGGREN, W.A., KENT, D.V., AUBRY, M-P., AND HARDENBOL, J., 1995,Geochronology, Time Scales and Global Stratigraphic Correlation:SEPM, Special Publication 54, 386 p.

BERGGREN, W.A., AND OUDA, K., 2003a, Upper Paleocene–lower Eoceneplanktonic foraminiferal biostratigraphy of the Dababiya section,Upper Nile Valley (Egypt): Micropaleontology, v. 49, supplement 1,p. 61–92.

BERGGREN, W.A., AND OUDA, K., 2003b, Upper Paleocene–lower Eoceneplanktonic foraminiferal biostratigraphy of the Qreiya (Gebel AbuHad) section, Upper Nile Valley (Egypt): Micropaleontology, v. 49,suppl. 1, p. 105–122.

BLAKEY, R.C., 2008, WEB presentation at: (http://jan.ucc.nau.edu/~rcb7/370_Devonian_2globes.jpg).

CRAMER, B.S., AND KENT, D.V., 2005, Bolide summer: the Paleocene/Eocene thermal maximum as a response to an extraterrestrial trig-ger: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 224, p.144–166.

CRICK, R.E., ELLWOOD, B.B., EL HASSANI, A., FEIST, R., AND HLADIL, J., 1997,Magneto Susceptibility event and cyclostratigraphy (MSEC) of theEifelian–Givetian GSSP and associated boundary sequences in NorthAfrica and Europe: Episodes, v. 20, p. 167–175.

DUPUIS, C., AUBRY, M., STEURBAUT, E., BERGGREN, W.A., OUDA, K.,MAGIONCALDA, R., CRAMER, B.S., KENT, D.V., SPEIJER, R.P., AND HEILMANN-CLAUSEN, C., 2003, The Dababiya Quarry Section: Lithostratigraphy,clay mineralogy, geochemistry and paleontology: Micropaleontol-ogy, v. 49, suppl. 1, p. 41–59.

ELLWOOD, B.B., HROUDA, F., AND WAGNER, J.-J., 1988, Symposia on mag-netic fabrics: Introductory comments: Physics of the Earth andPlanetary Interiors, v. 51, p. 249–252.

ELLWOOD, B.B., CRICK, R.E., AND EL HASSANI, A., 1999, The MagnetoSusceptibility event and cyclostratigraphy (MSEC) method used ingeological correlation of Devonian rocks from Anti-Atlas Morocco:American Association of Petroleum Geologists, Bulletin, v. 83, p.1119–1134.

ELLWOOD, B.B., CRICK, R.E., EL HASSANI, A., BENOIST, S.L., AND YOUNG, R.H.,2000, Magnetosusceptibility event and cyclostratigraphy methodapplied to marine rocks: Detrital input versus carbonate productiv-ity: Geology, v. 28, p. 1135–1138.

ELLWOOD, B.B., CRICK, R.E., FERNANDEZ, J.L.G., SOTO, F.M., TRUYOLS-MASSONI,M., EL HASSANI, A., AND KOVAS, E.J., 2001, Global correlation usingmagnetic susceptibility data from Lower Devonian rocks: Geology,v. 29, p. 583–586.

ELLWOOD, B.B., MACDONALD, W.D., WHEELER, C., AND BENOIST, S.L., 2003,The K-T boundary in Oman: identified using magnetic susceptibil-ity field measurements with geochemical confirmation: Earth andPlanetary Science Letters, v. 206, p. 529–540.

ELLWOOD, B.B., GARCÍA-ALCALDE, J.L., EL HASSANI, A., HLADIL, J., SOTO,F.M., TRUYÓLS-MASSONI, M., WEDDIGE, K., AND KOPTIKOVA, L., 2006,Stratigraphy of the Middle Devonian boundary: formal definition ofthe susceptibility magnetostratotype in Germany with comparisonsto sections in the Czech Republic, Morocco and Spain:Tectonophysics, v. 418, p. 31–49.

ELLWOOD, B.B., BRETT, C.E., AND MACDONALD, W.D., 2007, Magneto-susceptibility stratigraphy of the Upper Ordovician Kope Forma-tion, Northern Kentucky: Palaeogeography, Palaeoclimatology,Palaeoecology, v. 243, p. 42–54.

ELLWOOD, B.B., TOMKIN, J.H., RATCLIFFE, K.T., WRIGHT, M., AND KAFAFY,A.M., 2008, High resolution magnetic susceptibility and geochemis-try for the Cenomanian/Turonian boundary GSSP with correlationto time equivalent core: Palaeogeography, Palaeoclimatology,Palaeoecology, v. 261, p. 105–126, doi: 10.1016/j.palaeo.2008.01.005.

GRADSTEIN, F.M., OGG, J.G., AND SMITH, A.G., 2004, A Geologic Time Scale2004: Cambridge, U.K., Cambridge University Press, 589 p.

HARRINGTON, G.J., KEMP, S.J., AND KOCH, P.L., 2004, Palaeocene–Eoceneparatropical floral change in North America: responses to climate

FIG. 13.—Spectral character of the Zumaia MS data (Fig. 4). TheFourier transform of the detrended raw MS data (solid line) iscompared with the predicted location of ~ 100 kyr, ~ 52 kyr,~ 40 kyr, and ~ 18.7 kyr Milankovitch cyclicity (shadedvertical bars). The arrow represents the smoothed spline MScycles from bar logs in Figure 4.


change and plant migration: Geological Society of London, Journal,v. 161, p. 173–184.

HROUDA, F., 1994, A technique for the measurement of thermal changesof magnetic susceptibility of weakly magnetic rocks by the CS-2apparatus and the KLY-2 Kappabridge: Geophysical Journal Inter-national, v. 118, p. 604–612.

JENKINS, G.M., AND WATTS, D.G., 1968, Spectral Analysis and Its Applica-tions: San Francisco, Holden-Day, 525 p.

KENT, D.V., AND DUPUIS, C., 2003, Paleomagnetic study of the Paleocene–Eocene Tarawan Chalk and Esna Shale: Dual polarityremagnitizations of Cenozoic sediments in the Nile Valley (Egypt):Micropaleontology, v. 49, suppl. 1, p. 139–146.

KNOX, R.W.O’B., AUBRY, M., BERGGREN, W.A., DUPUIS, C., OUDA, K.,MAGIONCALDA, R., AND SOLIMAN, M., 2003, The Qreiya Section at GebelAbu Had: Lithostratigraphy, clay mineralogy, geochemistry andbiostratigraphy: Micropaleontology, v. 49, suppl. 1, p. 93–104.

LOURENS, L.J., SLUIJS, A., KROON, D., ZACHOS, J.C., THOMAS, E., RÖHL, U.,BOWLES, J., AND RAFFI, I., 2005, Astronomical pacing of late Palaeoceneto early Eocene global warming events: Nature, v. 435, p. 1083–-1087.

LU, G.Y., KELLER, G., ADATTE, T., ORTIZ, N., AND MOLINA, E., 1996, Long-term (10(5) or short-term (10(3) delta C-13 excursion near thePalaeocene–Eocene transition: Evidence from the Tethys: TerraNova, v. 8, p. 347–355.

LU, G.Y., ADATTE, T., KELLER, G., AND ORTIZ, N., 1998, Abrupt climatic,oceanographic and ecologic changes near the Paleocene–Eocenetransition in the deep Tethys basin: The Alamedilla section, south-ern Spain: Eclogae Geologicae Helvetiae, v. 91, p. 293–306.

NORRIS, R.D., AND RÖHL, U., 1999, Carbon cycling and chronology ofclimate warming during the Paleocene/Eocene transition: Nature,v. 401, p. 775–778.

NUNES, F., AND NORRIS, R.D., 2006, Abrupt reversal in ocean overturn-ing during the Palaeocene/Eocene warm period: Nature, v. 439, p.60–63.

OUDA, K., 2003, The Paleocene/Eocene boundary in Egypt: An over-view: Micropaleontology, v. 49, suppl. 1, p. 15–40.

OUDA, K., AND BERGGREN, W. A., 2003, Biostratigraphic correlation of theUpper Paleocene–Lower Eocene succession in the Upper Nile Val-ley: A synthesis: Micropaleontology, v. 49, suppl. 1, p. 179–212.

RÖHL, U., BRALOWER, T.J., NORRIS, R.D., AND WEFER, G., 2000, New chronol-ogy for the late Paleocene thermal maximum and its environmentalimplications: Geology, v. 28, p. 927–930.

SALVADOR, A., ed., 1994, International Stratigraphic Guide: A Guide toStratigraphic Classification, Terminology, and Procedure, SecondEdition: International Union of Geological Sciences (IUGS) and theGeological Society of America (GSA) co-publishers, 214 p.

SCHMITZ, B., PEUCKER-EHRENBRINK, B., HEILMANN-CLAUSEN, C., ABERG, G.,ASARO, F., AND LEE, C.-T. A., 2004, Basaltic explosive volcanism, butno comet impact, at the Paleocene–Eocene boundary: high-resolu-tion chemical and isotopic records from Egypt, Spain and Denmark:Earth and Planetary Science Letters, v. 225, p. 1–17.

SCOTESE, C.R., 1998, PALEOMAP Animations: PALEOMAP Project,Video Disk.

SOLIMAN, M.F., 2003, Chemostratigraphy of the Paleocene/Eocene (P/E)boundary sediments at Gabal el-Qreiya, Nile Valley, Egypt: Micro-paleontology, v. 49, suppl. 1, p. 123–138.

STOREY, M., DUNCAN, R.A., AND SWISHER, C.C., III, 2007, Paleocene–Eocenethermal maximum and the opening of the Northeast Atlantic: Sci-ence, v. 316, p. 587–589.

SWARTZENDRUBER, L.J., 1992, Properties, units and constants in magne-tism: Journal of Magnetic Materials, v. 100, p. 573–575.

THOMAS, D.J., AND BRALOWER, T.J., 2005, Sedimentary trace element con-straints on the role of North Atlantic Igneous Province volcanism inlate Paleocene–early Eocene environmental change: Marine Geol-ogy, v. 217, p. 233–254.

TRIPATI, A., AND ELDERFIELD, H., 2005, Deep-sea temperature and circula-tion changes at the Paleocene–Eocene Thermal Maximum: Science,v. 308, p. 1894–1898.

VAN DOBENECK, T., AND SCHMIEDER, F., 1999, Using rock magnetic proxyrecords for orbital tuning and extended time series analyses into thesuper- and sub-Milankovitch Bands, in Fischer, G., and Wefer, G.,eds., Uses of Proxies in Paleoceanography: Examples from the SouthAtlantic: Berlin, Springer-Verlag, p. 601–633.

WEAST, R.C., ed., 1982, CRC Handbook of Chemistry and Physics: BocaRaton, Florida, CRC Press, Inc.

WEEDON, G.P., JENKYNS, H.C., COE, A.L., AND HESSELBO, S.P., 1999, Astro-nomical calibration of the Jurassic time-scale from cyclostratigraphyin British mudrock formations: Royal Society of London, Philosophi-cal Transactions A, v. 357, p. 1787–1813.

WESTERHOLD, T., RÖHL, U., LASKER, J., RAFFI, I., BOWLES, J., LOURENS, L.J., AND

ZACHOS, J.C., 2007, On the duration of magnetochrons C24r and C25nand the timing of early Eocene global warming events: Implicationsfrom the Ocean Drilling Program Leg 208 Walvis Ridge depth transect:Paleoceanography, v. 22, PA2201, doi: 10.1029/2006PA001322.

WESTERHOLD, T., RÖHL, U., LASKER, J., RAFFI, I., FORNACIARI, E., MONECHI, S.,REALE, V., BOWLES, J., AND EVANS, H.F., 2008, Astronomical calibrationof the Paleocene time: Palaeogeography, Palaeoclimatology,Palaeoecology, v. 257, p. 377–403.

WING, S.L., BAO, H., AND KOCH, P.L., 2000, An early Eocene cool period?Evidence for continental cooling during the warmest part of theCenozoic, in Huber, B.T., MacLeod, K.G., and Wing, S.L., eds., WarmClimates in Earth History: Cambridge, U.K., Cambridge UniversityPress, p. 197–237.

WING, S.L., HARRINGTON, G.J., SMITH, F.A., BLOCH, J.I., BOYER, D.M., AND

FREEMAN, K.H., 2005, Transient floral change and rapid global warm-ing at the Paleocene–Eocene boundary: Science, v. 310, p. 993–996.

ZACHOS, J.C., PAGANI, M., SLOAN, L., THOMAS, E., AND BILLUPS, K., 2001,Trends, rhythms, and aberrations in global climate 65 Ma to present:Science, v. 292, p. 686–693.

ZACHOS, J.C., RÖHL, U., SCHELLENBERG, S.A., SLIUJS, A., HODELL, D.A., KELLY,D.C., THOMAS, E., NICOLO, M., RAFFI, I., LOURENS, L.J., MCCARREN, H., AND

KROON, D., 2005, Rapid acidification of the ocean during the Pale-ocene–Eocene Thermal Maximum: Science, v. 308, p. 1611–1615.


Documents

2010 Paleo-Eo.pdf