PALEOCEANOGRAPHY, VOL. 17, NO. 3, 10.1029/2001PA000623, 2002

Oceanic anoxic events and plankton evolution: Biotic response totectonic forcing during the mid-Cretaceous
R. Mark Leckie
Department of Geosciences, University of Massachusetts-Amherst, Amherst, Massachusetts, USA

Timothy J. Bralower
University of North Carolina at Chapel Hill, Chapel Hill, North Carolina, USA

Richard Cashman
Department of Geosciences, University of Massachusetts-Amherst, Amherst, Massachusetts, USA
Received 19 January 2001; revised 27 March 2002; accepted 27 March 2002; published 23 August 2002.

[1] Mid-Cretaceous (Barremian-Turonian) plankton preserved in deep-sea marl, organic-rich shale, and pelagic

carbonate hold an important record of how the marine biosphere responded to short- and long-term changes in
the ocean-climate system. Oceanic anoxic events (OAEs) were short-lived episodes of organic carbon burial thatare distinguished by their widespread distribution as discrete beds of black shale and/or pronounced carbon
isotopic excursions. OAE1a in the early Aptian (120.5 Ma) and OAE2 at the Cenomanian/Turonian boundary(93.5 Ma) were global in their distribution and associated with heightened marine productivity. OAE1b spans
the Aptian/Albian boundary (113–109 Ma) and represents a protracted interval of dysoxia with multiple
discrete black shales across parts of Tethys (including Mexico), while OAE1d developed across eastern andwestern Tethys and in other locales during the latest Albian (99.5 Ma). Mineralized plankton experienced
accelerated rates of speciation and extinction at or near the major Cretaceous OAEs, and strontium isotopic
evidence suggests a possible link to times of rapid oceanic plateau formation and/or increased rates of ridge crest
volcanism. Elevated levels of trace metals in OAE1a and OAE2 strata suggest that marine productivity may
have been facilitated by increased availability of dissolved iron. The association of plankton turnover and carbon
isotopic excursions with each of the major OAEs, despite the variable geographic distribution of black shaleaccumulation, points to widespread changes in the ocean-climate system. Ocean crust production andhydrothermal activity increased in the late Aptian. Faster spreading rates [and/or increased ridge length] drove along-term (Albian–early Turonian) rise in sea level and CO2-induced global warming. Changes in ocean
circulation, water column stratification, and nutrient partitioning lead to a reorganization of plankton community
structure and widespread carbonate (chalk) deposition during the Late Cretaceous. We conclude that there were
important linkages between submarine volcanism, plankton evolution, and the cycling of carbon through the
INDEX TERMS: 3030 Marine Geology and Geophysics: Micropaleontology; 4267 Oceanography: General:marine biosphere.
Paleoceanography; 4815 Oceanography: Biological and Chemical: Ecosystems, structure and dynamics; 4855 Oceanography: Biologicaland Chemical: Plankton; 9609 Information Related to Geologic Time: Mesozoic; KEYWORDS: mid-Cretaceous, oceanic anoxic events,
planktic foraminifera, calcareous nannofossils, radiolarians, submarine volcanism

1. Introduction[2] The mid-Cretaceous (124– 90 Ma) was a time oftransition in the nature of the ocean-climate system. Changes
were brought about by increased rates of tectonic activity
and shifting paleogeography [e.g., Larson and Pitman,
1972; Barron, 1987; Larson, 1991a, 1991b; Jones et al.,
1994; Ingram et al., 1994; Hay, 1995; Poulsen et al., 1999a,
2001; Jones and Jenkyns, 2001]. Intervals of the Early
Cretaceous were characterized by relatively cool high latitudes including evidence for ice rafting [Frakes and Francis, 1988; Weissert and Lini, 1991; Frakes et al., 1992; Stolland Schrag, 1996; Weissert et al., 1998; Clarke and Jenkyns,
Copyright 2002 by the American Geophysical Union.
0883-8305/02/2001PA000623$12.00

1999; Ferguson et al., 1999; Frakes, 1999]. However,
beginning in the Aptian, increased ocean crust production(greater rates of seafloor spreading and/or increased ridge
length), coupled with active midplate and plate margin
volcanism contributed to a greenhouse world of rising sea
level and warming global climate that peaked in the earlyTuronian but persisted through early Campanian time [Haysand Pitman, 1973; Kominz, 1984; Arthur et al., 1985; Rich etal., 1986; Larson, 1991a, 1991b; Huber et al., 1995, 2002;
Clarke and Jenkyns, 1999]. Elevated levels of volcanicallyderived CO2 in the atmosphere resulted in ice-free poles andcontributed to a weak meridional temperature gradient andan increasingly active hydrologic cycle [Barron and Washington, 1985; Barron et al., 1989, 1995; Huber et al., 1995,
2002; Schmidt and Mysak, 1996; Hay and DeConto, 1999;
Poulsen et al., 1999b]. The accompanying widespread burial

13 - 1

13 - 2LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

of organic carbon punctuated this long-term global warmth anumber of times with episodes of reverse greenhouse climate
marked by cooling and perhaps glaciation [e.g., Arthur et al.,
1988; Weissert and Lini, 1991; Weissert et al., 1998; Hochuli
et al., 1999; Kuypers et al., 1999; Stoll and Schrag, 2000].
Many researchers have suggested that ocean circulation was
driven in part by the sinking of warm, saline waters in
subtropical regions of excessive evaporation [Chamberlin,
1906; Brass et al., 1982; Southam et al., 1982; Wilde andBerry, 1982; Barron, 1983; Arthur et al., 1985, 1987; Barronand Peterson, 1990; Woo et al., 1992; Barron et al., 1993,
1995; Hay, 1995; Johnson et al., 1996; Schmidt and Mysak,1996]. However, ocean general circulation model experiments show that the southern high latitudes were likely thedominant sites of deep water formation during the midCretaceous with subtropical convection restricted to isolated
basins [Poulsen et al., 2001].
[3] At times during the mid-Cretaceous, dysoxic andanoxic conditions developed in oxygen minimum zones
along continental margins of the tropical Tethys Sea, in
restricted epicontinental seas, and in basins of the widening
North and South Atlantic Ocean basins. These conditions
led to the regional deposition of rhythmically bedded
sedimentary sequences including organic-rich black shale[e.g., Dean et al., 1978; McCave, 1979; Arthur and Premoli
Silva, 1982; de Boer, 1982; Cotillon and Rio, 1984]. The
observed cyclicity has been interpreted as orbitally forcedchanges in climate, which controlled biogenic carbonate
flux, productivity, and/or redox conditions at the seafloor[e.g., Arthur et al., 1984; de Boer and Wonders, 1984;
Fischer, 1986; Herbert and Fischer, 1986; Herbert et al.,
1986; Pratt and King, 1986; Gale et al., 1993; Arthur andSageman, 1994; Sageman et al., 1998]. Oceanic anoxic
events (OAEs), on the other hand, were generally shortlived (<1 Myr) episodes of organic carbon burial characterized by the widespread distribution of discrete beds ofblack shale and/or pronounced positive carbon isotopic
excursions (typically >1.5 – 2%) (Figure 1) [Schlanger
and Jenkyns, 1976; Arthur and Schlanger, 1979; Jenkyns,
1980; Arthur et al., 1987, 1990; Bralower et al., 1993]. The
major mid-Cretaceous OAEs are associated with the accumulation of marine organic matter [Erbacher et al., 1996].
[4] As sea level rose and global climate warmed during
the Albian-Turonian, there was a marked shift from organiccarbon-rich black shale deposition in the basins of western(Atlantic-Caribbean) and eastern (Mediterranean) Tethys tocarbonate (chalk) deposition along flooded continental
margins and in newly created or expanding epicontinental
seas [e.g., Arthur and Premoli Silva, 1982; Bréhéret et al.,
1986; Premoli Silva et al., 1989; Tornaghi et al., 1989]. Bylate Cenomanian-early Turonian time the deepening gateway between the basins of the North and South Atlantic
altered ocean circulation and improved deep water ventilation through much of Tethys [Arthur and Natland, 1979;
Tucholke and Vogt, 1979; Summerhayes, 1981; Arthur andPremoli Silva, 1982; de Graciansky et al., 1982; Cool,
1982; Zimmerman et al., 1987; Leckie, 1989; Arthur etal., 1990; Poulsen et al., 1999a, 2001].
[5] The mid-Cretaceous was also a time of rapid radiationand turnover in the marine plankton [Lipps, 1970; Haq,
1973; Tappan and Loeblich, 1973; Bujak and Williams,
1979; Caron and Homewood, 1983; Roth, 1987; Leckie,1989], benthic foraminifera [Sliter, 1977, 1980; Sikora andOlsson, 1991; Kaiho, 1999; Holbourn and Kuhnt, 2001],
molluscs [Vermeij, 1977; Theyer, 1983; Signor and Vermeij,
1994], and terrestrial plants [Hickey and Doyle, 1977;
Retallack and Dilcher, 1986; Lidgard and Crane, 1988;
Crane et al., 1995]. This evolutionary activity, part of theso-called ‘‘Mesozoic revolution’’ of Vermeij [1977], transformed many of these groups on an ocean-wide basis.
Diversity increased dramatically as organisms invaded
new habitats and partitioned the changing ecospace. The
reorganization of the global biosphere during the midCretaceous paralleled the marked changes in the oceanclimate system, suggesting a causal relationship between
biotic evolution and environmental stimuli [Fischer andArthur, 1977; Rich et al., 1986; Leckie, 1989; Thurow et al.,
1992; Vermeij, 1995].
[6] Planktic protists with mineralized skeletons of calcium
carbonate and silica, namely, the autotrophic calcareous
nannoplankton and heterotrophic planktic foraminifera andradiolarians, are generally abundant and well preserved in
mid-Cretaceous marine sediments. Radiolarians show highrates of evolutionary turnover (extinction plus radiation) ator near the OAEs [Erbacher et al., 1996; Erbacher andThurow, 1997], and detailed studies of individual events
have demonstrated that the calcareous nannoplankton andplanktic foraminifera were likewise influenced to varying
degrees by the OAEs [Hart, 1980; Leckie, 1985; Hart andBall, 1986; Bralower, 1988; Bralower et al., 1993, 1994;
Erba, 1994; Leckie et al., 1998; Huber et al., 1999; Premoli
Silva et al., 1999]. These data suggest that the OAEs were

Figure 1. (opposite) The mid-Cretaceous record of major black shales and oceanic anoxic events (OAEs) in the context ofthe carbon isotopic record [Erbacher et al., 1996; Bralower et al., 1999], changing global sea level [Haq et al., 1988], andseawater chemistry [Bralower et al., 1997]. See text for discussion of emplacement history of submarine large igneous
provinces (LIPs: Ontong-Java, Manihiki, and Kerguelen Plateaus and the Caribbean Plate). Short-term sea level changes areshown as the dark shaded line, and the long-term record of sea level is shown with the thick solid line (adapted from Haq etal. [1988]). OAE1a in the early Aptian (‘‘Selli event’’), OAE1b spanning the Aptian/Albian boundary (including the‘‘Jacob,’’ ‘‘Paquier,’’ and ‘‘Urbino’’ events), and OAE2 at the Cenomanian/Turonian boundary (‘‘Bonarelli event’’) eachcorrespond to negative 87Sr/86Sr excursions indicative of elevated submarine volcanism [Bralower et al., 1997; Larson and
Erba, 1999; Jones and Jenkyns, 2001]. Note that the 87Sr/86Sr ratio in marine carbonate declines to its most negative values
in the latest Aptian, roughly coincident with increased rates of ocean crust production [Kominz, 1984; Larson, 1991a,
1991b]. Initiation of increased spreading rates drove the long-term (Albian-Turonian) rise of global sea level. After rising to aplateau by the late Albian, perhaps due to increased rates of continental weathering with global warming, the 87Sr/86Sr ratio
once again dropped sharply at or near the time of OAE2.

LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

13 - 3

13 - 4LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

foci of evolutionary change for the plankton during the midCretaceous.
[7] Beginning in middle to late Aptian time (116– 113Ma), ocean crust production increased significantly [Kominz,
1984; Larson, 1991a, 1991b]; the timing is constrained bybiostratigraphic and strontium isotopic data [Bralower etal., 1997; Jones and Jenkyns, 2001]. In addition toincreased rates of ocean crust production during the midCretaceous, plumes of hot, buoyant rock rose through themantle forming anomalously thick and extensive oceanic
plateaus, termed large igneous provinces (LIPs). These
include (1) the Ontong-Java and Manihiki Plateaus andthe intervening Nova Canton Trough (125– 118 Ma withrenewed activity 96 – 84 Ma [Mahoney et al., 1993;
Tejada et al., 1996; Larson and Kincaid, 1996; Larson,
1997; Larson and Erba, 1999; Mahoney et al., 2002]), (2)the Kerguelen Plateau (KP) (116– 110 Ma for the southern KP and Rajmahal, 110– 108 Ma for Elan Bank, and95 –85 Ma for the central KP and Broken Ridge [Whitechurch et al., 1992; Coffin and Eldholm, 1994; Frey et al.,
1999; Pringle and Duncan, 2000; Shipboard ScientificParty, 2000]), and (3) the Caribbean Plate (94– 87 Ma[Sinton et al., 1998]). An important consequence of thissubmarine volcanism was increased hydrothermal activity,
particularly at the spreading centers, which caused secularchanges in ocean chemistry [Vogt, 1989; Hardie, 1996;
Bralower et al., 1997; Stanley and Hardie, 1998; Jonesand Jenkyns, 2001].
[8] A succession of environmental changes links intraplate, mantle plume volcanism with biotic turnover,enhanced productivity, and widespread burial of marineorganic matter during the early Aptian OAE (120.5 Ma[Bralower et al., 1994; Erba, 1994; Larson and Erba,
1999]). Others have further suggested that the OAE at the
Cenomanian/Turonian boundary (93.5 Ma) was likewise
related to submarine volcanism, including the formation of
the Caribbean Plateau [Schlanger et al., 1987; Orth et al.,
1993; Ingram et al., 1994; Sinton and Duncan, 1997; Kerr,
1998; Snow and Duncan, 2001]. The temporal associationof increased submarine volcanism with the early Aptian andCenomanian/Turonian boundary OAEs as well as additional
activity at the end of the Aptian is strongly supported bystrontium isotopic data [Bralower et al., 1997; Jones andJenkyns, 2001] (Figure 1).
[9] The purpose of this paper is to review a 33 million
year record of mid-Cretaceous plankton evolution in thecontext of the short-lived OAEs and the longer-term trendof rising sea level, changing paleogeography, and warming
global climate. Here we present new data on the evolutionary rates of calcareous plankton based on an integrated
calcareous plankton biostratigraphy and geochronology[Bralower et al., 1995, 1997; Erba et al., 1996]. Our
plankton-based analysis supports the hypothesis that submarine volcanism, namely, oceanic plateau formation,
coupled with increased ocean crust production and hydrothermal activity was an important catalyst of marine
productivity and black shale deposition during the major
mid-Cretaceous OAEs. However, the early Aptian world
was very different from that of Cenomanian/Turonianboundary time, and therefore the individual OAEs sharesome similarities as well as important differences. In
particular, the Albian stage (112.2 – 98.9 Ma) records
changes in the foci of organic carbon burial and pelagic
carbonate deposition as well as major changes in planktic
foraminiferal diversity, size, and wall structure [e.g., Bréhéret et al., 1986; Leckie, 1989; Premoli Silva and Sliter,
1999]. The degree to which these patterns were controlled
by tectonically driven changes in ocean circulation, watercolumn structure, productivity, and plankton community
dynamics and to what extent such changes may or may not
be related to the spatial and temporal distribution of the
OAEs are central to the paleoceanographic research questions considered here.

2. Oceanic Anoxic Events[10] There were arguably between two and seven OAEs
during the mid-Cretaceous [Schlanger and Jenkyns, 1976;
Arthur and Schlanger, 1979; Jenkyns, 1980; Arthur et al.,
1987; Sliter, 1989a; Arthur et al., 1990; Bralower et al.,
1993, 1999; Erbacher et al., 1996; Erbacher and Thurow,1997]. Erbacher et al. [1996] distinguished between OAEs
associated with rising sea level, heightened marine productivity (type II kerogen), and positive d13C excursions, socalled POAEs (P for productivity), and DOAEs (D for
detrital) characterized by falling sea level and terrestrialorganic matter (type III kerogen). The two most widespread
OAEs are the early Aptian OAE1a (Livello Selli; 120.5
Ma) and the Cenomanian-Turonian boundary OAE2 (Livello Bonarelli; 93.5 Ma) (Figure 1); these are both POAEs
according to Erbacher et al. [1996]. The concentration oforganic matter in distinct, widely distributed beds of blackshale was facilitated by reduced terrigenous sedimentation
during transgression and/or the incursion of upwellinginduced oxygen minima across the upper slope and shelfwith rising sea level [Hallam and Bradshaw, 1979; Arthur
et al., 1987; Schlanger et al., 1987; Loutit et al., 1988;
Arthur et al., 1990; Bralower et al., 1993; Arthur andSageman, 1994]. However, major unresolved questions
remain: what could sustain widespread and elevated marine
productivity, and were the individual OAEs triggered bysimilar forcing mechanisms?
2.1. OAE1a (Early Aptian)[11] The characteristics of the individual OAEs illustrate
their complex nature. The onset of OAE1a (Selli event), forexample, is preceded by a sharp negative d13C excursion(0.5 – 3.0%), an abrupt decrease in 87Sr/86Sr isotope values(Figure 1), increased trace metal concentrations, and amajor demise in the nannoconids (the ‘‘nannoconid crisis’’[Erba, 1994]), followed by an abrupt and prolonged positive d13C excursion (>2%) and short-lived (0.5 – 1.0 Myr.)black shale deposition [Sliter, 1989a; Bralower et al., 1994,
1997, 1999; Föllmi et al., 1994; Jenkyns, 1995; Menegatti
et al., 1998; Jones and Jenkyns, 2001]. The series of events
across the Globigerinelloides blowi-Leupoldina cabri biozone boundary has been linked to the Ontong-Java Pacific‘‘superplume’’ eruption [Larson, 1991a, 1991b; Tarduno etal., 1991; Erba, 1994; Larson and Erba, 1999]. The initialLECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

negative d13C excursion is recorded in marine carbonates at
Resolution Guyot in the Pacific [Jenkyns, 1995], in Alpine
sections of northern Tethys [Menegatti et al., 1998], and in
Mexico [Bralower et al., 1999] as well as in fossil woodfrom the Isle of Wight [Gröcke et al., 1999]. Hypotheses toaccount for the negative excursion include (1) increased
upwelling rates of nutrient- and 12C-rich intermediatewaters [Menegatti et al., 1998], (2) accelerated hydrologic
cycle and higher weathering rates induced by CO2-driven
global warming [Menegatti et al., 1998], (3) isotopicallylight, mantle-derived CO2 associated with the eruption of
Ontong Java-Manihiki prior to an interval of enhanced
productivity and organic carbon burial [Bralower et al.,
1994; Larson and Erba, 1999], and (4) dissociation of
isotopically light methane hydrates in continental margin
sediments [Jahren and Arens, 1998; Opdyke et al., 1999;
Jahren et al., 2001]. Larson and Erba [1999] hypothesizedthat iron fertilization associated with the superplume eruption triggered increased productivity during OAE1a.
2.2. OAE2 (Cenomanian/Turonian Boundary Interval)[12] The OAE2 (Bonarelli event) also displays a marked
positive d13C excursion (>2%) related to the widespread
burial of marine organic matter [Scholle and Arthur, 1980;
Summerhayes, 1981, 1987; Pratt and Threlkeld, 1984;
Arthur et al., 1987, 1990; Schlanger et al., 1987; Jarvis etal., 1988; Hilbrecht et al., 1992; Thurow et al., 1992; Galeet al., 1993; Pratt et al., 1993; Jenkyns et al., 1994;
Sugarman et al., 1999], and a marked drop in 87Sr/86Sr
isotopic values has been attributed to submarine volcanism[Ingram et al., 1994; Bralower et al., 1997; Jones andJenkyns, 2001] (Figure 1). Elevated rates of seafloor spreading and subduction are implicated by the volume, thickness,and extent of ashfall deposits (bentonites) through the
Cenomanian/Turonian boundary interval in the WesternInterior Sea of North America [e.g., Kauffman, 1984;
Kauffman and Caldwell, 1993]. Additional supporting evidence for increased hydrothermal activity comes fromelevated levels of trace metals in marls and organic-richmudrocks of the southern part of the Western Interior Sea as
tropical water masses invaded the seaway with rising sea
level [Orth et al., 1993; Leckie et al., 1998; Snow andDuncan, 2001].
[13] A number of researchers have suggested that OAE2
was likewise triggered by iron fertilization associated withmantle plume volcanism [Sinton and Duncan, 1997; Kerr,
1998]. However, most of the dated lava flows from theyounger volcanic sequence on Ontong-Java Plateau as well
as flows from the Caribbean Plate are too young (92 – 88Ma), and volcanism on the Kerguelen Plateau was subaerial
by late Cenomanian-early Turonian time [Shipboard Scientific Party, 2000]. Alternatively, OAE2 may have been
triggered by the injection of warm saline intermediate ordeep waters that created favorable conditions for the verticaladvection of nutrients, widespread productivity, expansionof oxygen minima, and the accumulation of organic matter.
For example, Huber et al. [1999, 2002] showed that middle
bathyal waters (500 – 1000 m) in the western North
Atlantic abruptly warmed from 15 to 20C in the latest
Cenomanian. These paleotemperature estimates are the

13 - 5

warmest known Cretaceous or Cenozoic intermediatewaters. The loss of water column density gradients at theend of the Cenomanian may account for the extinction ofdeeper-dwelling foraminifera (Rotalipora spp., Globigerinelloides bentonensis), a condition that would have also
facilitated the vertical advection of nutrients to fuel primary
productivity [Leckie et al., 1998; Huber et al., 1999].
[14] Arthur et al. [1987] suggested that the creation ofwarm saline deep waters in the expanding epicontinental
seas drove upwelling during OAE2. In addition, an abruptchange in Atlantic deep water circulation created by thebreaching of the deep water sill separating the North andSouth Atlantic Ocean basins [Arthur and Natland, 1979;
Tucholke and Vogt, 1979; Summerhayes, 1981, 1987; Cool,
1982; Zimmerman et al., 1987; Poulsen et al., 1999a, 2001]
may have also facilitated nutrient delivery to the euphotic
zone by ventilating the deep North Atlantic and adjacent
Tethyan margins. Thus the high productivity of OAE2 may
have been triggered by changes in deep and intermediatewater circulation and/or source(s) of water mass production,
which was sustained by a volcanically spiked, nutrient-richwater column. The accumulation and preservation oforganic matter was further enhanced by the widespreadlatest Cenomanian-early Turonian transgression and by thewarm, oxygen-poor intermediate and deep waters. Burial oforganic carbon during OAE2 was modulated by orbitally
forced climate cyclicity [Gale et al., 1993; Sageman et al.,1998]. For example, data from central Tunisia indicates thatthe accumulation of organic matter during OAE2 wasdriven by the precessional cycle (20 kyr) for the durationof a single 400 kyr eccentricity cycle [Caron et al., 1999].
2.3. Other OAEs and Possible OAEs
[15] Additional intervals of black shale accumulation
include the late Aptian (116 Ma), latest Aptian – early
Albian (OAE1b; 113– 109 Ma), late Albian (OAE1c and
OAE1d; 102 and 99.2 Ma, respectively), and midCenomanian (96 Ma [Arthur et al., 1990; Bralower etal., 1993, 1999; Bréhéret, 1994; Erbacher et al., 1996,
2001; Wilson and Norris, 2001]). A black shale event in thelate Aptian, between OAE1a and OAE1b in the Globigerinelloides algerianus biozone [Bralower et al., 1999; Sliter,
1999], may prove to be an OAE. It is represented by adiscrete black shale bed in the Calera Limestone of northern
California (paleolatitude of 15 – 17N in the eastern
Pacific [Sliter, 1999; Premoli Silva and Sliter, 1999]) andin northeastern Mexico where an accompanying large
negative-to-positive excursion in d13C, similar to OAE1a,
implies a widespread distribution (Figure 1) [Bralower etal., 1999]. This interval of the late Aptian has been linked tocooling and an eustatic sea level fall [Weissert and Lini,
1991; Weissert et al., 1998].
[16] The multiple black shales of OAE1b (upper Ticinella
bejaouaensis and Hedbergella planispira biozones) aremostly restricted to Mexico and the North Atlantic basin(western Tethys) and the Mediterranean (eastern Tethys)
region [e.g., Arthur and Premoli Silva, 1982; Bréhéret et al.,
1986; Premoli Silva et al., 1989; Bralower et al., 1993,1999]. This interval is associated with cooling and sea level
fall in the latest Aptian and subsequent sea level rise during

13 - 6LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

the early Albian [Weissert and Lini, 1991; Weissert et al.,1998]. OAE1c in the lower upper Albian Biticinella breggiensis biozone has been identified in central Italy, the U.S.
western interior, and Australia (‘‘Toolebuc’’) and is associated with the accumulation of terrigenous organic matter[Pratt and King, 1986; Bralower et al., 1993; Coccioni andGaleotti, 1993; Haig and Lynch, 1993; Erbacher et al.,1996].
[17] OAE1d (Rotalipora appenninica biozone), on theother hand, is widely preserved as a black shale across
Tethys (‘‘Breistroffer’’) with sporadic occurrences in theSouth Atlantic, southern Indian, and eastern Pacific Ocean
basins; it is associated with marine organic matter [Bréhéretand Delamette, 1989; Bréhéret, 1994; Erbacher et al., 1996;
Wilson and Norris, 2001]. Cyclic black shales in the uppermost Albian of Ocean Drilling Program (ODP) Site 1052 on
Blake Nose are correlative to OAE1d and correspond to aninterval marked by collapse of upper water column stratification caused by intensified winter mixing and reduced
summer stratification [Wilson and Norris, 2001]. Alternatively, the isotopic data of the thermocline-dwelling genus
Rotalipora suggest that warming of intermediate waters
may have also contributed to the collapse of upper watercolumn density gradients. Another possible OAE occurs in
the mid-Cenomanian of Tethys where it is associated with apositive carbon excursion (1%), marine organic matter,and an extinction event in the radiolaria [Erbacher et al.,
1996; Stoll and Schrag, 2000]. Noteworthy are the large(1 – 2%) carbon isotopic excursions [e.g., Weissert andLini, 1991; Weissert et al., 1998; Bralower et al., 1999; Stolland Schrag, 2000; Wilson and Norris, 2001] and radiolarian
turnover events [e.g., Erbacher et al., 1996] that are concentrated at or near the times of both OAE1b and OAE1d,suggesting widespread environmental changes.
2.4. Significance of OAE1b (Latest Aptian – Early
Albian)[18] OAE1b is particularly important in the analysis that
follows because it represents a major transition in the natureof mid-Cretaceous tectonics, sea level, climate, lithofacies,
and marine plankton communities. This is an interval that
contains several prominent black shales spanning the timeof the Aptian/Albian boundary [Arthur and Premoli Silva,
1982; Bréhéret et al., 1986; Bréhéret and Delamette, 1989;
Premoli Silva et al., 1989; Tornaghi et al., 1989; Bréhéret,
1991, 1994; Erbacher et al., 1996, 1998, 1999; Erbacher
and Thurow, 1997]. A sea level fall near the Aptian/Albianboundary [Weissert and Lini, 1991; Bréhéret, 1994; Weissert

et al., 1998] separates uppermost Aptian black shales(‘‘Jacob’’ in the Vocontian Basin, and ‘‘113’’ in central
Italy) from lower Albian black shales (‘‘Paquier’’ and‘‘Leenhardt’’ events in the Vocontian Basin, and the ‘‘Monte
Nerone’’ and ‘‘Urbino’’ events in the Apennines of central
Italy; Figure 1). Erbacher et al. [1998] interpreted theuppermost Aptian Jacob event of the Vocontian Basin as a
DOAE (formed by detrital input; ‘‘fed black shale’’) and thelower Albian Paquier and Leenhardt events as POAEs(driven by productivity and rising sea level; ‘‘condensed
black shale’’).
[19] The middle to late Aptian interval leading up to
OAE1b is characterized by heavy d13Ccarb and d13Corg
values, punctuated by two episodes of sharply lighter
isotopic ratios (Figure 1). Specifically, d13C values continue to rise after the early Aptian OAE1a, suggesting thecontinued widespread burial of organic carbon through themid-Aptian (Leupoldina cabri and Globigerinelloides ferreolensis biozones), in parallel with rising sea level[Menegatti et al., 1998]. This is followed by depleted
values in the middle-late Aptian (Globigerinelloides algerianus biozone), enriched values again in the late Aptian(Hedbergella trocoidea and lower Ticinella bejaouaensis
biozones), and depleted values in the latest Aptian (upper
T. bejaouaensis biozone) before rising again in the early
Albian (H. planispira zone [Weissert and Lini, 1991;
Erbacher et al., 1996; Weissert et al., 1998; Bralower etal., 1999]). Weissert and Lini [1991] attribute the negative
carbon excursions to global cooling, ice sheet growth, andsea level fall in the late Aptian and near the Aptian-Albianboundary as a positive feedback to the prolonged episodesof Corg burial in the L. cabri-G. ferreolensis biozones andagain in the H. trocoidea-T. bejaouaensis biozones. The
occurrence of ice-rafted debris and cool temperaturesduring the Aptian and Aptian/Albian transition in Australia support an interpretation of high-latitude glaciation(s) [Frakes and Francis, 1988; Ferguson et al.,
1999; Frakes, 1999].
[20] Biogenic data from Deep Sea Drilling Program(DSDP) Site 545 [Leckie, 1984, 1987] suggest that theupper Aptian – basal Albian interval off northwest Africa
was characterized by high productivity based on the abundance of radiolarians and benthic organisms (siliceoussponges, benthic foraminifera, echinoids, and ostracodes);the latter responded favorably to the enhanced export
production [Leckie, 1987; Berger and Diester-Haas, 1988;
Herguera and Berger, 1991] (Figure 2). The sharp drop inpercent planktic foraminifera corresponds with the extinc-

Figure 2. (opposite) High biological productivity, represented in this expanded record of OAE1b at DSDP Site 545 off
Morocco as suggested by the abundance of radiolarians and benthic organisms [Leckie, 1984, unpublished data]. Increased
export production supported the rich benthic ecosystems [Leckie, 1987; Berger and Diester-Haas, 1988; Herguera andBerger, 1991]. Black shale lithofacies are weakly developed at this site, although elevated levels of organic carbon arepreserved in cores 545-42 and 545-43 [Hinz et al., 1984]. The lithofacies and biofacies of Site 545 bear strong resemblance
to those of the European margin of Tethys, particularly the Piobbico core from the Umbria-Marche Basin of central Italy[Premoli Silva et al., 1989; Tornaghi et al., 1989] (cores 545-55 to 47 are equivalent to unit 17 to the lower part of unit 14
in the Piobbico core; cores 545-46 to 545-41 are equivalent to the upper part of unit 14 to unit 11; strata above thedisconformity in Site 545 are equivalent to midunit 2 in the Piobbico core). The larger planktic foraminifera were wiped out
in the OAE1b interval, and tiny species dominate the assemblages. This was a watershed event in the evolutionary historyof planktic foraminifera (see text).

Cenom.

Stage

lower

upper

l.

AlbianAptianL. acut.
(NC11)

Calc.
NannosA. alb.
(NC9)

P. columnata (NC8)

R. angustus (NC7)

R. reicheli

Planktic
Forams

260

320

33
34

43

500
520

52
53
54
55
56

G. algerian.

G. ferreolen.

unzoned

540

460
480

LO P. chen.

H. trocoidea

440

420

400

380

360

47
48
49
50
51

46

44
45

41
42

38
39
40

340

300

32

35
36
37

280

30
31

29

0

20

40
60
Percent

80

100"Jacob"

OAE 1b"Paquier"(to total forams)

28% Planktic ForamsCore% Recovery

H. planisp.

R. ticinen.

R. appen.

R. brotzeni

T. bejaouaensis

E. turriseiffelii(NC10)

NC8B

NC8A

NC7C

NC7B

upper

Biostratigraphy

20

¥

¥

¥

40
60
Percent

radiolarians

¥

¥

¥

80

sponge spicules

calcispheres

Moroccan Margin

DSDP Site 545(to total biogenic component)

0

¥

¥

¥

¥

5
Percent

•

•

•

•

10 0

ostracods

echinoderms

fish bone(to total biogenics)% Sponge Spicules, % Ostracods,
Radiolarians, &Echinoderms,
Calcispheres& Fish••••% Glauconite

20

•

•
•

•
•

•

•

40
60
Percent

•

•

•

•

•

•

OAE1d

80

(to total mineral component)

100

•
•

LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

13 - 7

13 - 8LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

tion of Planomalina cheniourensis, the earliest planktic
foraminifer with a peripheral keel, and the pseudoextinctionof Nannoconus truittii [Leckie, 1984; this study]. The low
planktic:benthic ratio (percent planktics) across the Aptian/
Albian boundary at Site 545 may be partially related toincreased carbonate dissolution, although preservation of
the tiny planktic and benthic species is moderately good togood with minimal test fragmentation and etching of thecalcareous tests. However, correlative strata from the Apennines of central Italy show pronounced dissolution in thebasal Albian [Premoli Silva et al., 1989; Tornaghi et al.,
1989; Erba, 1992]. Despite the presumed high productivity
off Morocco, black shale development is weak at thislocality. High sedimentation rates (>35 m/Myr) diluted thelithologic expression of the correlative Jacob event at Site
545. In contrast, the basal Albian strata accumulated atsignificantly reduced rates (<15 m/Myr (Leckie, unpublished data)) and the presence of a thin dark-colored layer
with elevated total organic carbon in core 545-42 [Hinz etal., 1984; Bralower et al., 1993] likely correlates with thePaquier black shale event (Figure 2). Like the Vocontian
Basin of southeast France and the Umbria-Marche Basin ofcentral Italy, the onset of OAE1b off northwest Africa ismarked by an abrupt change in lithology and biota in theuppermost Aptian, just below the level of the Jacob and 113
black shale events.
[21] Erbacher et al. [2001] provide planktic and benthic
foraminiferal stable isotopic data across the basal Albianpart of OAE1b (correlative to the Paquier level) at ODP Site
1049 in the western tropical North Atlantic (Blake Nose)that show a sharp increase in planktic-benthic d18O gradients across the black shale event. These authors attributethe findings to increased stratification of the water column
by surface water warming and/or increased runoff andsuggest that this black shale formed as a megasapropel byanalogy to Plio-Pleistocene sapropel accumulation in the
Mediterranean. Kuypers et al. [2001] concluded that severe
oxygen depletion affected the water column during thisevent on the basis of the dominant fraction of organic matter
derived from chemoautotrophic Crenarchaeota bacteria inthis black shale. However, the basal Albian Paquier andlower Albian Leenhardt events in the Vocontian Basin ofsoutheast France have been attributed to elevated primary
productivity [Bréhéret, 1994; Erbacher et al., 1998, 1999],
rather than increased thermohaline stratification. Erbacher
et al. [1996] show three rock eval analyses for the intervalaround OAE1b; two have type II (marine) kerogen, and one
has type III (terrestrial). These findings suggest that multiple
triggers, including productivity, sea level, or climaticallydriven organic carbon burial events, characterize the broadinterval of OAE1b.
[22] A marked lithologic change from poorly oxidized
sediments with discrete black shale(s) to highly oxidized
sediments is observed in lower Albian strata of ODP Site
1049 in the western North Atlantic (Blake Nose [Norris etal., 1998]) and at DSDP Site 511 in the southern South
Atlantic (Falkland Plateau [Ludwig et al., 1983]), suggesting the possibility of an intermediate water (<1000 m)
connection between the North and South Atlantic Ocean
basins by early Albian time. In addition, the unconformity

separating the basal Albian and upper Albian strata at Site
545 (Figure 2) may be related to the inferred increase inintermediate water ventilation and circulation, thereby terminating OAE1b at middle (to lower?) bathyal depths. Poor
ventilation at greater depths in the eastern North Atlantic
persisted through much of the Albian [e.g., Summerhayes,
1981, 1987].

3. Methods[23] One of the biggest challenges in studying the evolutionary history of ancient organisms is building the chronology required to calculate rates of change. Recent studiesof mid-Cretaceous calcareous nannofossils and planktic
foraminifera from deep-sea sections have established an
integrated calcareous plankton biostratigraphy and improved geochronology [Bralower et al., 1993, 1994, 1995,
1997; Erba et al., 1996] (Figure 3). In this study, a 33 Myrinterval encompassing the late Barremian to late Turonian(123– 90 Ma) was divided into 1 million year increments
using the timescale of Gradstein et al. [1994] and biostratigraphic age model of Bralower et al. [1997]. The stratigraphic ranges of 91 planktic foraminifera and 78
calcareous nannofossil taxa were compiled from numerous
sources, including land sections and deep-sea sites (Figures
4 and 5). New and/or revised planktic foraminiferal range
data from DSDP Sites 545 and 547 on the Moroccan
continental margin [Leckie, 1984], DSDP Sites 364 and511 in the South Atlantic, and ODP Site 763 on the
Exmouth Plateau off northwest Australia (R. Cashman,
M.S. thesis in progress, University of Massechusetts,
2002) are integrated with published data. The resulting
compilations are based on a conservative taxonomic framework and best estimates of biostratigraphic range. Evolutionary rates were calculated following the method of Wei
and Kennett [1986] (Tables 1 and 2).
[24] Here we compare the nature of biotic change as
expressed by evolutionary rates in the planktic foraminifera
and calcareous nannofossils (this study) and radiolarians[Erbacher et al., 1996; Erbacher and Thurow, 1997] withthe temporal and spatial distribution of the OAEs and otherproxies of global change, including strontium isotopic ratios(87Sr/86Sr) [Bralower et al., 1997], carbon isotopes (d13Ccarb
and d13Corg [Erbacher et al., 1996; Bralower et al., 1999]),
and oxygen isotopic analyses (d18O) of individual species ofplanktic foraminifera and monogeneric benthic foraminifera[Norris and Wilson, 1998; this study]. We also examined
trends in planktic foraminiferal size and morphology as
proxies of plankton trophic structure and upper watercolumn structure, respectively.
[25] Isotope paleoecology (d18O and d13C) of individual
species of planktic foraminifera is a useful tool to assess
depth habitat or presence of photosymbionts in ancient
species [e.g., Berger et al., 1978; Fairbanks and Wiebe,
1980; Fairbanks et al., 1982; Deuser et al., 1981; Gasperi
and Kennett, 1992; Ravelo and Fairbanks, 1992, 1995;
D’Hondt and Arthur, 1995; Norris, 1996, 1998; Pearson,
1998; Pearson et al., 2001]. Interplanktic species isotopic
gradients reveal the nature of the upper water column(mixed layer and upper thermocline where most plankticLECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

Figure 3. Integrated calcareous plankton biostratigraphy used in this study (modified from Bralower et al.
[1997]) following revised range of Leupoldina cabri from Erba et al. [1999] and Premoli Silva et al. [1999].

13 - 9Figure 4. Interpreted stratigraphic ranges of 91 species of planktic foraminfera used to calculate the evolutionary rates
presented in this study (Table 1).

13 - 10LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTIONLECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

13 - 11

Figure 5. Interpreted stratigraphic ranges of 78 species of calcareous nannofossils used to calculate theevolutionary rates presented in this study (Table 2). There are 28 species that range through the entire
mid-Cretaceous interval as represented by the wide shaded bar.

foraminifera live), including aspects of density structure andproductivity, while planktic-benthic isotopic gradients serve
as proxies of water column structure, productivity, andbottom water age. Changing patterns in mid-Cretaceousplanktic foraminiferal morphological and trophic diversity
have been linked with upper water column stratification andproductivity [Sliter, 1972; Hart, 1980; Caron and Homewood, 1983; Leckie, 1987, 1989; Premoli Silva and Sliter,
1999].
[26] Here we present new data on the isotope paleoecology of latest Aptian and middle Albian species of planktic
foraminifera. We measured the carbon and oxygen isotopesof multiple species of well-preserved planktic foraminifera
and a single epifaunal benthic taxon (Gavelinella) from four
uppermost Aptian and middle Albian samples of DSDP Site
392A on the Blake Nose, western subtropical North Atlantic(Table 3) [see Gradstein, 1978, plates]. The samples arefrom the type of clay-rich hemipelagic sediments shown toyield reliable paleotemperature estimates from foraminiferal
calcite [e.g., Norris and Wilson, 1998; Pearson et al., 2001;
Wilson and Norris, 2001]. Analyses were conducted atWoods Hole Oceanographic Institution on a Finnigan
MAT 252 with a precision of better than 0.03% for d13C
and 0.08% for d18O. Our results are integrated with other13 - 12LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

Table 1. Planktic Foraminiferal Evolutionary Data Based on theStratigraphic Ranges Presented in Figure 4a

Table 2. Calcareous Nannofossil Evolutionary Data Based on theStratigraphic Ranges Presented in Figure 5aInterval,
Ma

Age

S

FO

LO

rs

re

rd

rt

Interval,
Ma

Age

S

FO

LO

rs

re

rd

rt

90 – 91
91 – 92
92 – 93
93 – 94
94 – 95
95 – 96
96 – 97
97 – 98
98 – 99
99 – 100
100 – 101
101 – 102
102 – 103
103 – 104
104 – 105
105 – 106
106 – 107
107 – 108108 – 109
109 – 110
110 – 111
111 – 112
112 – 113
113 – 114
114 – 115
115 – 116
116 – 117
117 – 118
118 – 119119 – 120
120 – 121
121 – 122
122 – 123

90.5
91.5
92.5
93.5
94.5
95.5
96.5
97.5
98.5
99.5
100.5
101.5
102.5
103.5
104.5
105.5
106.5
107.5
108.5
109.5
110.5
111.5
112.5
113.5
114.5
115.5
116.5
117.5
118.5
119.5
120.5
121.5
122.5

26
24
23
25
23
19
22
19
19
23
22
22
15
12
12
11
8
7
6
7
5
6
13
11
16
16
16
15
20
22
18
14
12

2
3
3
4
5
2
4
1
2
7
3
7
3
0
1
3
1
1
1
2
0
1
3
0
2
0
3
1
4
4
4
2
0

5
0
2
5
2
2
5
1
1
6
6
3
0
0
0
0
0
0
0
2
0
1
9
0
5
2
0
2
6
6
0
0
0

0.08
0.13
0.13
0.16
0.22
0.11
0.18
0.05
0.11
0.30
0.14
0.32
0.20
0.00
0.08
0.27
0.13
0.14
0.17
0.29
0.00
0.17
0.23
0.00
0.13
0.00
0.19
0.07
0.20
0.18
0.22
0.14
0.00

0.19
0.00
0.09
0.20
0.09
0.11
0.23
0.05
0.05
0.26
0.27
0.14
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.29
0.00
0.17
0.69
0.00
0.31
0.13
0.00
0.13
0.30
0.27
0.00
0.00
0.00

0.12
0.13
0.04
0.04
0.13
0.00
0.05
0.00
0.05
0.04
0.14
0.18
0.20
0.00
0.08
0.27
0.13
0.14
0.17
0.00
0.00
0.00
0.46
0.00
0.19
0.13
0.19
0.07
0.10
0.09
0.22
0.14
0.00

0.27
0.13
0.22
0.36
0.30
0.21
0.41
0.11
0.16
0.57
0.41
0.45
0.20
0.00
0.08
0.27
0.13
0.14
0.17
0.57
0.00
0.33
0.92
0.00
0.44
0.13
0.19
0.20
0.50
0.45
0.22
0.14
0.00

90 – 91
91 – 92
92 – 93
93 – 94
94 – 95
95 – 96
96 – 97
97 – 98
98 – 99
99 – 100
100 – 101
101 – 102
102 – 103
103 – 104
104 – 105
105 – 106
106 – 107
107 – 108108 – 109
109 – 110
110 – 111
111 – 112
112 – 113
113 – 114
114 – 115
115 – 116
116 – 117
117 – 118
118 – 119119 – 120
120 – 121
121 – 122
122 – 123

90.5
91.5
92.5
93.5
94.5
95.5
96.5
97.5
98.5
99.5
100.5
101.5
102.5
103.5
104.5
105.5
106.5
107.5
108.5
109.5
110.5
111.5
112.5
113.5
114.5
115.5
116.5
117.5
118.5
119.5
120.5
121.5
122.5

60
60
63
67
63
62
62
59
58
59
59
59
58
55
56
56
55
52
50
50
50
49
48
46
46
45
44
45
45
44
44
41
38

0
1
2
4
0
0
3
1
0
0
0
2
3
0
1
1
3
2
0
0
0
2
2
0
1
1
0
0
3
1
3
3
2

0
0
4
6
0
0
0
0
0
1
0
0
1
0
1
1
0
0
0
0
0
0
0
0
0
0
0
1
0
3
0
0
0

0.00
0.02
0.03
0.06
0.00
0.00
0.05
0.02
0.00
0.00
0.00
0.03
0.05
0.00
0.02
0.02
0.05
0.04
0.00
0.00
0.00
0.04
0.04
0.00
0.02
0.02
0.00
0.00
0.07
0.02
0.07
0.07
0.05

0.00
0.00
0.06
0.09
0.00
0.00
0.00
0.00
0.00
0.02
0.00
0.00
0.02
0.00
0.02
0.02
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.02
0.00
0.07
0.00
0.00
0.00

0.00
0.02
0.03
0.03
0.00
0.00
0.05
0.02
0.00
0.02
0.00
0.03
0.03
0.00
0.00
0.00
0.05
0.04
0.00
0.00
0.00
0.04
0.04
0.00
0.02
0.02
0.00
0.02
0.07
0.05
0.07
0.07
0.05

0.00
0.02
0.10
0.15
0.00
0.00
0.05
0.02
0.00
0.02
0.00
0.03
0.07
0.00
0.04
0.04
0.05
0.04
0.00
0.00
0.00
0.04
0.04
0.00
0.02
0.02
0.00
0.02
0.07
0.09
0.07
0.07
0.05

a
The data are plotted in Figure 6 and summarized in Figures 7 and 11. S,
species richness (or simple diversity); FO, number of first occurrences per
million year interval; LO, number of last occurrences per million year
interval; rs, rate of speciation ((1/S)FO); re, rate of extinction ((1/S)LO); rd,rate of diversification (rsre); rt, rate of turnover (rs + re). Rate ofspeciation (rs) = 1/species richness (S)  number of first occurrences (FO).
Rate of extinction (re) = 1/species richness (S)  number of last
occurrences (LO); rs and re are per species rates of speciation andextinction (1/S). Rate of diversification (rd) = rate of speciation (rs) rateof extinction (re). Rate of turnover (rt) = rate of speciation (rs) + rate ofextinction (re).

published data from Blake Nose DSDP and ODP sites[Norris and Wilson, 1998; Fassell and Bralower, 1999;
Huber et al., 1999].
[27] Planktic foraminiferal size data are based on thelargest dimension of all species known from a single
tropical-subtropical planktic foraminiferal biozone or composite of two biozones. These data are derived from scanning electron photomicrograph images from various
sources: late Barremian and early Aptian data are fromODP Site 641 on the Iberian margin (Leckie, unpublished)
and outcrop localities in Mexico [Longoria, 1974] andSpain [Coccioni and Premoli Silva, 1994] and late Aptian,early Albian and late Albian data are from DSDP Sites 545
and 547 on the Moroccan margin [Leckie, 1984] with
supplementary data from Longoria [1974]. Mid-Cretaceousplanktic foraminiferal size data are compared with extant
species [Hemleben et al., 1989] to illustrate the significant
differences in size and trophic specialization between thea
See Table 1 for explanation. The data are plotted in Figure 6 andsummarized in Figures 7 and 11. Rate of speciation (rs) = 1/species richness(S)  number of first occurrences (FO). Rate of extinction (re) = 1/species
richness (S)  number of last occurrences (LO); rs and re are per species
rates of speciation and extinction (1/S). Rate of diversification (rd) = rate ofspeciation (rs)rate of extinction (re). Rate of turnover (rt) = rate ofspeciation (rs) + rate of extinction (re).

Table 3. Stable Isotopic Data for Individual Species of Planktic
Foraminifera and the Benthic Foraminiferal Genus Gavelinella
From Four Samples from DSDP Hole 392A (Blake Nose)a
Sample Number

Taxon

Age

d13C

d18O

392A-2-1, 73 cm
392A-2-1, 73 cm
392A-2-1, 73 cm
392A-2-1, 73 cm
392A-3-1, 94 cm
392A-3-1, 94 cm
392A-3-1, 94 cm
392A-3-2, 102 cm
392A-3-2, 102 m
392A-3-2, 102 cm
392A-3-2, 102 cm
392A-3-2, 102 cm
392A-3-3, 73 cm
392A-3-3, 73 cm
392A-3-3, 73 cm
392A-3-3, 73 cm
392A-3-3, 73 cm

T. primula
H. planispira
Hedbergella n.sp.
Gavelinella sp.
T. primula
Hedbergella n.sp.
Gavelinella sp.
T. bejaouaensis
H. trocoidea
Blefuscuiana sp.
Gavelinella sp.
G. aptiense
T. bejaouaensis
Blefuscuiana sp.
P. cheniourensis
Gavelinella sp.
G. aptiense

105.2
105.2
105.2
105.2
106.2
106.2
106.2
112.5
112.5
112.5
112.5
112.5
112.8
112.8
112.8
112.8
112.8

2.202
2.202
2.052
1.338
2.341
2.048
1.286
4.016
4.058
4.155
3.511
4.068
3.783
3.897
3.684
3.193
3.757

2.455
2.059
2.007
0.371
2.150
1.847
0.374
0.255
0.257
0.415
0.779
0.835
0.096
0.247
0.525
0.566
0.721

a

The data are presented in Figures 9 and 10. Here n.sp. is new species.

LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

13 - 13

Figure 6. Evolutionary history of planktic foraminifera and calcareous nannofossils through the midCretaceous (Tables 1 and 2). Evolutionary rates calculated following the method of Wei and Kennett[1986]: (1) rate of speciation (rs) = 1/species richness (S)  number of first occurrences (FO) per million
year interval, (2) rate of extinction (re) = 1/species richness (S)  number of last occurrences (LO) per
million year interval, (3) rate of diversification (rd) = rs re, and (4) rate of turnover (rt) = rs + re. The‘‘nannoconid crisis’’ represents a global perturbation in calcareous nannoplankton communities that
produced a dramatic reduction in nannoconid carbonate production immediately preceding OAE 1a[Erba, 1994; Larson and Erba, 1999]. Nannoconids make a resurgence of abundance in the late Aptian
called the ‘‘Nannoconus truittii acme,’’ named for the dominant taxon [Mutterlose, 1989; Erba, 1994].
The N. truittii acme corresponds with the range of the earliest planktic foraminiferal taxon with aperipheral keel, Planomalina cheniourensis. Nannoconids abruptly drop in abundance at the extinctionlevel of P. cheniourensis in the latest Aptian [Erba, 1994; Bralower et al., 1994; this study]. We refer tothis second nannoconid crash in abundance as the ‘‘nannoconid minicrisis.’’ These two events mark theonset of OAE1b just prior to the ‘‘Jacob’’ black shale.

ancient and modern taxa. Plankton size in part reflects
trophic position within the microbial loop and/or grazing
food chain (see section 4).

4. Results
4.1. Evolutionary Rates of Mid-Cretaceous
Calcareous Plankton[28] The pattern of evolutionary activity in the calcareous
nannoplankton reveal the greatest rates of turnover (speciation plus extinction) in the early Aptian and at the

Cenomanian/Turonian boundary (Figure 6). These twointervals also correspond with the highest rates of extinction. There were typically 2 – 3 times more species ofcalcareous nannoplankton than planktic foraminifera
through much of the mid-Cretaceous, with the early to
middle Albian showing the greatest disparity between thetwo groups. Calcareous nannofossil diversity gradually rises
from 40 species in the late Barremian to a high of nearly
70 species in the early Turonian. The rate of diversification(speciation minus extinction) is greatest in the late Barremian and in the early late Aptian; other times of significant

13 - 14LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTIONLECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

diversification include the Aptian/Albian boundary interval,
middle Albian, late Albian, and mid-Cenomanian.
[29] Planktic foraminiferal evolution during the mid-Cretaceous was characterized by episodes of elevated rates ofturnover alternating with times of relatively diminishedrates (note the difference in scale between the planktic
foraminifera and the calcareous nannofossils). The highest
rates of turnover occurred in the early to middle Aptian,across the Aptian/Albian boundary, in the latest Albian, themid-Cenomanian, and at the Cenomanian/Turonian boundary (Figure 6). Particularly noteworthy are the three peaks
in turnover through the latest Aptian-early Albian interval,
with the greatest turnover of planktic foraminifera of theentire mid-Cretaceous occurring near the Aptian/Albianboundary (e.g., 69% of the late Aptian species became
extinct during the 112 –113 Ma interval, while there was a
23% speciation rate for a total of 92% species turnover).
This turnover resulted in the net loss of species and thelowest simple diversity of the study interval.
[30] In general, the Aptian was dominated by negative
rates of planktic foraminiferal diversification, with exceptions at the beginning of the stage and in the early late
Aptian. Positive rates characterize nearly all of the Albianwith the exception of the end of the Albian. The resulting
pattern of simple diversity through time shows a pattern ofincreasing species richness during the late Barremian thatpeaked in the early Aptian, decreased somewhat in the
middle Aptian, and remained relatively stable through muchof the late Aptian (Figure 6). There was a fourfold increasein diversity between the early and late Albian with total
species richness remaining relatively stable through the
mid-Turonian despite episodes of elevated turnover.
[31] An important result of this analysis is the strikingly
similar patterns of evolutionary turnover in the planktic
foraminifera and the radiolarians. Erbacher et al. [1996]
showed that the highest rates of turnover in the radiolaria
coincide with the major OAEs (Figure 7). Foraminiferal
evolutionary patterns clearly show that the highest rates of
speciation and/or extinction occur at or near the OAEs. For
the calcareous nannofossils the highest rates of turnoveroccur at OAE1a in the early Aptian and OAE2 at the
Cenomanian/Turonian boundary. In addition to the major
OAEs, additional perturbations in the ocean-climate system
may have occurred during the mid-Aptian, late Albian, andthe mid-Cenomanian on the basis of accelerated evolutionary activity in both the planktic foraminifera and thecalcareous nannoplankton.
4.2. Planktic Foraminiferal Size and Trophic Ecology[32] Planktic foraminiferal size varied significantly during the mid-Cretaceous (Figure 8). During the late Barremian, all species were <250 mm with an average size of
168 mm. Average size steadily increased toward the late

13 - 15

Aptian in parallel with increasing diversity (Figures 4 and6) and morphological variety and complexity [Leckie,1989]. The average size of planktic foraminifera reached
330 mm by the late Aptian with the largest taxon (Globigerinelloides algerianus) attaining maximum diameters of700 mm. Planktic foraminiferal size and diversity bothcrashed across the Aptian/Albian boundary; at this time theaverage size dropped to 230 mm. Average test size steadilyrose to 400 mm by the end of the Albian in parallel with thetrend in diversity. Cenomanian and Turonian planktic
foraminifera are similar to late Albian taxa in their overall
range of size.
[33] By comparison with extant species of planktic foraminifera the mid-Cretaceous planktic foraminifera are muchsmaller (Figure 8). Note the wide range of size and implied
trophic diversity of modern taxa (e.g., the average adultmaximum diameter, on a spine-free basis, of the 44 extant
species is 500 mm; range is 100 – 2000 mm). Nearly half
(19 of 44) of all modern planktic foraminifera bear longspines that provide infrastructure to support the sticky
rhizopodial network used to snare particulate organic matteror prey and/or to deploy photosymbionts. In addition, themodern taxa show a diverse array of feeding strategies,
many of which were probably unavailable to mid-Cretaceous taxa, including a variety of photosymbioses andcarnivory of microcrustaceans such as copepods [Hemleben
et al., 1989; Laybourn-Parry, 1992] (Figure 8). For example, ultrastructural evidence indicates that none of theseancient species had spines, and isotopic data suggest that no
mid-Cretaceous species contained photosymbionts [Norris
and Wilson, 1998], although very few species have been
analyzed with the appropriate isotopic technique to criticallyevaluate for photosymbionts [e.g., D’Hondt and Arthur,
1995; Norris, 1996, 1998].
4.3. Planktic Foraminiferal Isotope Paleoecology
and Upper Water Column Stratification[34] Several studies have utilized isotopic analyses ofindividual mid-Cretaceous planktic foraminiferal species
for the purposes of assessing paleoecology, water column
structure, and/or productivity [e.g., Corfield et al., 1990;
Huber et al., 1995, 1999, 2002; Norris and Wilson, 1998;
Fassell and Bralower, 1999; Erbacher et al., 2001]. Our
results show that latest Aptian age planktic foraminifera(Ticinella bejaouaensis zone) were probably weakly
stratified in the upper water column on the basis of a
0.58 – 0.62% difference between the lightest (Ticinella
bejaouaensis) and heaviest (Globigerinelloides aptiense)
d18O values (Figures 9g and 9h and Table 3). The
planktic-benthic isotopic gradients also show that thewater column was weakly stratified with deep/intermediate water temperatures of 10 – 11C and salinity corrected surface water temperatures (SSTs) of 14– 17C.

Figure 7. (opposite) Plankton evolutionary events relative to the OAEs showing percentage of species first appearances(speciation) or disappearances (extinction) through 1 million year increments. Radiolarian data are from Erbacher andThurow [1997]. Planktic foraminiferal and calcareous nannofossil extinction and speciation data are based on this study(Figure 6). For planktic foraminifera, larger triangles represent 19% change, and smaller triangles represent 15– 19% forcalcareous nannofossils, larger triangles represent >4% change, and smaller triangles represent 3 – 4%. Note that the highest
rates of plankton turnover (extinction plus speciation) occur at or near the major OAEs.

13 - 16

LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

Figure 8. Comparison of mid-Cretaceous and modern planktic foraminiferal size and trophic strategies.
Data for modern planktic foraminifera are from Hemleben et al. [1989]; mid-Cretaceous data are from
Longoria [1974], Leckie [1984], Coccioni and Premoli Silva [1994], and R. M. Leckie, unpublished datafrom ODP Hole 641C). Data are also presented on Figure 10.

LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

13 - 17

Cibicidoides

Figure 9. Stable isotopic data for multispecies mid-Cretaceous planktic foraminifera from Blake Nosein the western North Atlantic. (a – d) Data from the upper Albian-basal Cenomanian of ODP Hole 1052E(data are from Norris and Wilson [1998]; multiple data points for each species represent different size
fractions). (e – h) New data from the uppermost Aptian and middle Albian of DSDP Site 392A (Table 3).
By late middle Albian time (upper Ticinella primula zone),salinity-corrected SSTs had risen to 25– 28C, and deep/
intermediate water temperatures in the western North Atlantic had risen to 14 – 17C (middle Albian data from thesame site by Fassell and Bralower [1999] show benthic
temperatures as low as 12.5C and salinity-corrected SSTs
as low as 22C). While upper water column isotopic
gradients remained weak (0.30– 0.45%), the surface to deep
water gradient of 1.78 – 2.08% indicates that the watercolumn was significantly more stratified at this location
by middle Albian time (Figures 9e and 9f).

[35] When integrated with other published records fromthe Blake Nose in the western North Atlantic [Norris andWilson, 1998; Fassell and Bralower, 1999; Huber et al.,1999], the isotopic data show clear evidence for increasing
stratification of the upper water column during latest Aptianthrough Albian time in parallel with increasing planktic
foraminiferal diversity and test size (Figure 10). Not only
did deep water temperature rise during the Albian, but boththe planktic-benthic and interplanktic species d18O gradients(0.46 – 1.79 and 1.67 – 2.95%, respectively [Norris andWilson, 1998]) increased with rising sea level and rising13 - 18

LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

Figure 10. Paleotemperature estimates based on benthic (solid circles) and multispecies planktic (opencircles) foraminiferal data from the Blake Nose in the western North Atlantic [Norris and Wilson, 1998;
Fassell and Bralower, 1999; Huber et al., 1999; this study]. Planktic foraminiferal paleotemperature
estimates have been corrected for salinity (values in parentheses) following Zachos et al. [1994].
Calculations are based on the equation of Erez and Luz [1983] assuming dwater = 1.0%smow for anonglacial world [Shackleton and Kennett, 1975]. Note the following: (1) the decrease in planktic
foraminiferal test size and diversity across the Aptian/Albian boundary is associated with a fundamental
change in foraminiferal ultrastructure, (2) planktic foraminiferal simple diversity (species richness) andaverage test size generally track the pattern of long-term rising sea level and warming during the Albian(species of Cenomanian and Turonian planktic foraminifera are very similar to late Albian species in testsize), (3) relatively cool surface and deep water temperatures during the latest Barremian and latest
Aptian give way to a warming trend in both deep water and surface waters through the Albian, (4) the
sharp increase in d18Oplanktic gradient in the late Albian is associated with a marked increase in planktic
foraminiferal diversity, (5) increased water column stratification by the late Albian marks the onset ofwidespread deposition of calcareous ooze (chalk) with rising sea level, and (6) latest Aptian to earlyTuronian sea surface temperatures track the long-term trend of global sea level.

LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

sea surface temperatures (Figures 9a – 9d), indicating thatthe water column in the western North Atlantic became wellstratified by early late Albian time. Sea surface temperatures
peaked in the late Albian with values exceeding 30C
[Norris and Wilson, 1998]. Water column stratification
persisted through the Cenomanian-Turonian [Norris andWilson, 1998; Huber et al., 1999].

5. Discussion
5.1. The Evolutionary Record of Mid-Cretaceous
Plankton[36] Calcareous nannoplankton and radiolarians have arich history back to the Triassic and earlier. By contrast, theplanktic foraminifera experienced their first major radiations
from tiny ancestors (with test diameters <100 mm) during
the Early Cretaceous [Tappan and Loeblich, 1973; Bannerand Desai, 1988; Sliter, 1989b; Coccioni et al., 1992;
Coccioni and Premoli Silva, 1994; BouDagher-Fadel etal., 1997; Premoli Silva and Sliter, 1999]. The evolutionary
patterns of the planktic foraminifera and other plankton
show that this period of overall diversification was punctuated by episodes of extinction and speciation; the highest
rates of turnover occurred at or near the OAEs (Figure 7). In
the planktic foraminifera a reversal in the general trend ofincreasing size and morphological variety and complexity
occurred during latest Aptian-early Albian OAE1b when the
average shell diameter decreased sharply to <250 mm anddiversity plummeted (Figure 10).
[37] Although relatively little is known about the ecologyof mid-Cretaceous planktic foraminifera, we can makeinferences about trophic strategy on the basis of the sizeof these ancient species. In the modern plankton, only thesmallest protists (nannoplankton: 2 – 20 mm; microplankton:
20 –200 mm) benefit energetically by consuming bacteria(picoplankton: 0.2– 2 mm). For example, autotrophic nannoplankton and picoplankton (including cyanobacteria) andheterotrophic bacteria are consumed by heterotrophic flagellates, which in turn are consumed by ciliates and small
sarcodines [Azam et al., 1983; Laybourn-Parry, 1992;
Jumars, 1993; Rivkin et al., 1996]. The tiny test size ofplanktic foraminifera in the Early Cretaceous and during
OAE1a and OAE1b may in large part reflect limited trophic
strategies and the availability of food resources, withbacterivory as a likely feeding strategy in the smallest taxa.
Many species of small benthic foraminifera are known tograze on heterotrophic bacteria living on and within thesediments [Lee, 1980; Lipps, 1983; Goldstein, 1999]. Small
planktic foraminifera, by analogy, are also likely to exploit
heterotrophic bacteria, aerobic photoheterotrophic bacteria[Fenchel, 2001; Kolber et al., 2001], and/or cyanobacteria
by suspension feeding or by adapting a pseudobenthic
lifestyle on marine snow or other organic-rich flocs, which
are heavily colonized by bacteria [Azam, 1998; Azam andLong, 2001]. Other important food sources of these earlyplanktic foraminifera likely included phytoplankton andflagellates such as the bacterivorous choanoflagellates(2 – 10 mm in size [Laybourn-Parry, 1992; Jumars, 1993]).
[38] The nannoconids, an extinct group of heavily calcified calcareous nannoplankton, were the dominant pelagic

13 - 19

carbonate producers of the Early Cretaceous. However, thisgroup experienced a two-stage perturbation during theAptian (Figure 6). The early Aptian ‘‘nannoconid crisis’’
was the more dramatic of the two events and was associatedwith the emplacement of LIPs in the Pacific as a precursor
to OAE1a [Erba, 1994; Larson and Erba, 1999]. We
hypothesize that the Pacific superplume eruption in theearly Aptian resulted in reduced carbonate availability(CO23 ) associated with increased CO2 and lowered pH.
Accordingly, this created a deleterious effect on the heavier
calcifying forms like the nannoconids in addition to altering
nutricline, thermocline, and plankton ecosystem dynamics
as suggested by Erba [1994].
[39] The first occurrence of the aberrant planktic foraminiferal species Leupoldina cabri, with its distinctive clublike chambers, is nearly coincident with the ‘‘nannoconid
crisis’’ and the onset of OAE1a [Premoli Silva et al., 1999].
Leupoldinids dominate the low-abundance, low-diversity
planktic foraminiferal assemblages during the Selli eventin Italy. It has been suggested that taxa with radiallyelongate chambers, such as L. cabri, may have been adapted
to low-oxygen conditions in the upper water column [BouDagher-Fadel et al., 1997; Premoli Silva et al., 1999].
Planktic foraminifera experienced a marked turnover inthe early Aptian immediately following OAE1a (Figure 7),
which was followed by increasing test size and morphological complexity during the middle to late Aptian [Leckie,
1989; Premoli Silva and Sliter, 1999]. A cooler climate in
the early late Aptian (Globigerinelloides algerianus biozone[Weissert and Lini, 1991; Weissert et al., 1998]) may havecreated cooler deep waters and stronger water column
density gradients thereby facilitating the diversification ofplanktic foraminifera at this time [e.g., Caron and Homewood, 1983; Leckie, 1989; Premoli Silva and Sliter, 1999].
[40] Toward the end of the Aptian the total range of
Planomalina cheniourensis, the first keeled planktic foraminifer, coincides with a nannoconid resurgence referred toas the ‘‘Nannoconus truittii acme’’ [Erba, 1994] (Figure 6).
The widespread distribution of the N. truittii assemblage has
been attributed to warming surface waters [Mutterlose,
1989]. A second nannoconid ‘‘minicrisis’’ coincides withthe extinction level of P. cheniourensis during the latest
Aptian (within the Ticinella bejaouaensis biozone) marking
the onset of falling sea level [Bréhéret, 1994], global cooling [Weissert and Lini, 1991; Weissert et al., 1998], and a
protracted interval of organic carbon burial (OAE1b) thatspans the Aptian/Albian boundary (Figure 1). Increased
ocean crust production near the end of the Aptian andduring the early Albian may have affected planktic ecosystems in a manner similar to OAE1a, including the loss of theheaviest calcified nannofossils and planktic foraminifera.
The relatively cool paleotemperature estimates for this timeinterval may also reflect lower seawater pH associated withelevated levels of submarine volcanism and increasing
atmospheric pCO2 [Zeebe, 2001]. Planktic foraminifera
suffered their greatest rates of extinction and turnover acrossthe Aptian/Albian boundary (Figure 11). The radiolarians
were likewise affected while the calcareous nannoplankton,other than the nannoconids, were largely unaffected. Foremost was a major evolutionary change in the ultrastructure13 - 20LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

Figure 11. Summary of the major geochemical, tectonic, sea level, and plankton evolutionary events
associated with the mid-Cretaceous oceanic anoxic events. Note the concentration of evolutionary
turnover events (speciation plus extinction) with the OAEs. Also note that OAE1a, OAE1b, and OAE2 are
temporally associated with increased submarine volcanic activity as indicated by the lower 87Sr/86Sr ratios[Bralower et al., 1997], and all three are linked to increased burial of marine organic matter [e.g., Arthur etal., 1987; Erbacher et al., 1996; Larson and Erba, 1999]. We hypothesize that submarine volcanism andhydrothermal activity at the spreading centers helped to fuel the elevated levels of marine productivity
during the OAEs by way of iron fertilization of the water column. In addition to submarine volcanism thestrontium isotope record is also influenced by continental weathering and runoff [e.g., Jones and Jenkyns,
2001], and the rise in 87Sr/86Sr ratios during the Albian may in large measure record increased weathering
rates with rising sea level and global warming despite the elevated level of ocean crust production thatsustained high global sea level through much of the Late Cretaceous. The high productivity associatedwith OAE1d may have been facilitated by an ocean already preconditioned by dissolved iron.

of planktic foraminiferal tests. This was a watershed eventin planktic foraminiferal evolution as the trochospirally
coiled taxa with small pores, many with distinctive poremounds, were replaced by taxa with larger pores andsmooth walls (Figure 10).
[41] During middle and late Albian time the planktic
foraminifera diversified rapidly, particularly during the late
Albian when keeled morphologies reappeared [Leckie,
1989; Premoli Silva and Sliter, 1999] (Figures 6 and 10).
There was a pronounced increase in the degree of calcification in many taxa between the middle and late Albian aswell as a continued increase in size, diversity, and morphological complexity. The development of muricae (spikes ofcalcite) on species of the genera Hedbergella, Rotalipora,and Praeglobotruncana, rugae (rows of calcite) on H.
libyca, and peripheral keels, raised sutures, and/or adumbilical ridges on species of Rotalipora, Praeglobotruncana,and Planomalina illustrate the propensity for calcification
by late Albian time. This pattern of evolutionary change inthe planktic foraminifera tracks the overall trend of rising
sea level and global warming. Stable isotopic evidence
points to exploitation of the uppermost water column by
way of depth stratification in the mixed layer and along thethermocline as density gradients within the upper ocean,and between the surface and deep ocean increased [Norris
and Wilson, 1998] (Figures 9 and 10). A major phase ofturnover affecting the planktic foraminifera and radiolarians
coincides with OAE1d in the latest Albian. The collapse of

LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

density gradients during OAE1d [Wilson and Norris, 2001]
may account for increased rates of extinction and speciationin both of these groups.
[42] The productivity event of OAE2 (Cenomanian/Turonian boundary) had a major effect on the radiolarians andcalcareous nannoplankton. The latter group experienced
their largest turnover of the mid-Cretaceous [Bralower,
1988] (Figure 11). Deeper-dwelling planktic foraminifera
were the most severely affected by OAE2 as evidenced by
the loss of the genus Rotalipora, possibly due to the
expansion of the oxygen minimum zone into thermocline
waters [Hart, 1980; Wonders, 1980; Leckie, 1985, 1989].
Alternatively, abrupt warming of deep waters, to nearly
20C at middle bathyal depths, may have caused a breakdown of the thermocline thereby effectively eliminating theniche of the deeper-dwelling species [Huber et al., 1999].
The extinction of deep-sea benthic foraminifera across the
Cenomanian/Turonian boundary [Kaiho et al., 1993; Kaihoand Hasegawa, 1994; Kaiho, 1998] may be related to thiswarming event analogous to the deep-sea benthic foraminiferal extinction associated with the Paleocene-Eocene thermal maximum [Tjalsma and Lohmann, 1983; Kennett andStott, 1991; Thomas and Shackleton, 1996].
5.2. Submarine Volcanism[43] Vermeij [1995] hypothesized that massive submarine
volcanism and hydrothermal activity coupled with rising sea
level and CO2-induced global warming were the principal
triggers to the major biosphere-scale revolutions of thePhanerozoic, including the rise of mineralized planktic
protists during the Mesozoic. He argued that when raw
materials and energy (especially temperature) become available to organisms at unusually high rates, opportunities forevolutionary innovation and diversification are enhanced
through increased access to nutrients and adaptation.
Nutrients are made available to the marine plankton throughcontinental weathering, upwelling of deep and/or intermediate waters, and volcanism. Biological activity, including
nutrient recycling, is strongly temperature-dependent [Brasier, 1995; Vermeij, 1995]. Therefore trends in global temperature and sea level tend to run parallel with marine
productivity and diversification. The mid-Cretaceous was
one of those times, and submarine volcanism was anintegral part of plankton resource economics through its
influence on increased CO2, global temperatures, andnutrient supply to the ocean [Vogt, 1989].
[44] The 87Sr/86Sr ratio of seawater decreased markedly
during three intervals of the mid-Cretaceous corresponding
with OAE1a, OAE1b, and OAE2 [Bralower et al., 1997;
Jones and Jenkyns, 2001] (Figure 11). These authorsinterpreted the lowered (less radiogenic) values as episodesof effusive emplacement of oceanic plateaus (mantle
plumes) or hydrothermal activity associated with increased
ocean crust production. Other studies have proposed important linkages between increased submarine volcanism(hydrothermal activity), primary productivity, and blackshale deposition [Vogt, 1989; Bralower et al., 1994; Erba,
1994; Ingram et al., 1994; Sinton and Duncan, 1997; Kerr,
1998; Larson and Erba, 1999]. Hydrothermal vents associated with spreading centers are major sources of dissolved

13 - 21iron and other biolimiting metals to the ocean and minor
sources of phosphorus [Froelich et al., 1982; Vermeij,
1995]. Horizontal and vertical advection of limiting
nutrients and trace elements, particularly iron, may haveprovided important catalysts for marine productivity during
times of elevated rates of seafloor spreading and submarine
volcanism, particularly for locations or times of weak watercolumn stratification.
[45 ] Large-scale ocean thermohaline circulation andwind-driven upwelling play a major role in supplying
nutrients to the euphotic zone of the world ocean. Welldeveloped density stratification in the low latitudes of themodern ocean coupled with seasonal stratification in the
midlatitudes greatly limits the vertical advection ofnutrients. These conditions create relatively focused zones
of elevated primary productivity in coastal and high-latitude
waters and in narrow bands of wind-driven oceanic divergence. In contrast, the reduced equator-to-pole temperature
gradients of the mid-Cretaceous ocean created generally
weaker water column density gradients. Such conditions
provided an opportunity for greater rates of vertical advection of nutrients at times of increased water mass production. For example, mesoscale eddies [Falkowski et al., 1991;
McGillicuddy and Robinson, 1997; McGillicuddy et al.,
1998; Williams and Follows, 1998; Oschlies and Garcon,
1998; Siegel et al., 1999], driven in part by the sinking ofwarm, saline waters, may have induced intermittent upwelling and lateral advection of nutrients in the upper layers ofthe mid-Cretaceous ocean. Planetary Rossby waves are also
effective mechanisms for the delivery of nutrients to thesurface ocean [Siegel, 2001; Uz et al., 2001]. Hydrothermal
megaplumes generated at the spreading centers and oceanic
plateaus may have also contributed to decreased watercolumn stratification and nutrient injection to the euphotic
zone [Vogt, 1989; Baker et al., 1995; Palmer and Ernst,
1998].
[46] Rates of seafloor spreading were 40– 50% greater
during the mid-Cretaceous than they are today [Richter etal., 1992]. Earlier studies have suggested that spreading
rates accelerated during the Aptian [Kominz, 1984; Larson,
1991a, 1991b]. The low 87Sr/86Sr isotopic ratios of the late
Aptian may mark the onset of increased ridge crest volcanism (Figure 11). Accelerated rates of seafloor spreading and/or increased ridge length during the latest Aptian initiatedthe long-term Albian-Turonian eustatic rise of sea level[Kominz, 1984; Haq et al., 1988; Larson, 1991a, 1991b]
and global warming [Huber et al., 1995; Clarke andJenkyns, 1999; Ferguson et al., 1999; Frakes, 1999].
According to the hypothesis of Stanley and Hardie[1998], increased hydrothermal activity associated withfaster spreading rates reduced oceanic Mg/Ca and preconditioned the ocean for the spread of pelagic carbonate on
continental margins and in epicontinental seas with rising
sea level.
5.3. Iron Fertilization and Sustained Productivity
During OAEs[47] Cyanobacteria contribute substantially to primary
productivity in nutrient-poor, oligotrophic waters [Choand Azam, 1988; Laybourn-Parry, 1992; Falkowski et al.,

13 - 22

LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

1998], and colonies of cyanobacteria support diverse
assemblages of consortial organisms, including protists[e.g., Glover et al., 1988; Capone et al., 1997]. Iron
availability greatly enhances nitrogen fixation by cyanobacteria [Codispoti, 1997; Falkowski, 1997], and it stimulates
increased production in the eukaryotic phytoplankton [Martin and Fitzwater, 1988; Martin et al., 1994; Coale et al.,1996a, 1996b; Frost, 1996]. The heterotrophic bacteria ofthe microbial loop compete with prokaryotic and eukaryotic
autotrophs for this limiting micronutrient [Kirchman, 1996].
Iron facilitates greater efficiency of carbon metabolism in
the heterotrophic bacteria; when iron is available, greaterbacterial biomass is produced and available for consumption by bacteriovores and higher trophic levels of thegrazing food chain [Azam et al., 1983; Azam, 1998]. Wheniron is limited, more of the dissolved organic carbon
consumed by the bacteria is respired to CO2 [Kirchman,
1996; Tortell et al., 1996].
[48] In the modern ocean, it is the larger phytoplankton
cells, principally the diatoms, that show the greatest productivity response to increased iron availability [Martin etal., 1994; Coale et al., 1996a, 1996b; Longhurst, 1996].
However, diatoms did not reach their pinnacle of importance in marine food webs until the Cenozoic [Harwood
and Nikolaev, 1995]. One of the few diatomaceous deposits
known from the mid-Cretaceous occurs near Antarctica andis of Aptian/Albian boundary age [Gersonde and Harwood,
1990; Harwood and Gersonde, 1990], a time roughly
coincident with OAE1b, onset of Rajmahal-Kerguelen Plateau volcanism, and higher rates of seafloor spreading(Figure 11).
[49] Iron limitation is widespread in the ocean today,
including the subtropical gyres as well as high-nutrient,
low-chlorophyll regions [Falkowski et al., 1998; Behrenfeld
and Kolber, 1999]. Mechanisms that enhance iron availability, such as increased aeolian dust deposition during theLast Glacial Maximum [Martin, 1990] or, as we and otherssuspect, iron fertilization by submarine volcanism during
the mid-Cretaceous, can result in increased rates of primary
productivity. However, in a warm Cretaceous world, elevated rates of denitrification in low-oxygen environments
may have limited the availability of fixed nitrogen requiredby the eukaryotic autotrophs [Codispoti, 1997; Falkowski,
1997], or phosphate may have become limited by elevatedabundances of nitrogen-fixing cyanobacteria [Codispoti,
1989; Toggweiler, 1999; Tyrell, 1999].
[50] We suggest that the availability of iron at times ofheightened submarine volcanism and oceanic crustal production stimulated and sustained the production of nitrogenfixing cyanobacteria and improved the efficiency of the
microbial loop, both of which facilitated greater export
production during the OAEs. Elevated trace metal abundances in OAE1a and OAE2 strata suggest a linkage
between these OAEs and submarine volcanism [Orth etal., 1993; Sinton and Duncan, 1997; Kerr, 1998; Larson
and Erba, 1999]. Of these two events it is possible that only
OAE1a was directly triggered by submarine volcanism andiron fertilization as suggested by Larson and Erba [1999].
[51] We hypothesize that increased ocean crust production
and hydrothermal activity beginning in the late Aptian(Figure 11) were indirectly responsible for stimulating andsustaining widespread production of marine organic matter
during the broad OAE1b event. By contrast, the subsequent
OAEs were likely initiated by changing forcing factors. For
example, beginning in the Albian, an increasingly stratified
water column limited the availability of nutrients to theoceanic photic zone, thereby altering planktic ecosystem
structure, reducing export production, and providing a
positive feedback for the global warming that persisted
through much of the Late Cretaceous. OAE1d (latest
Albian) and OAE2 were likely triggered by ocean-climate
events that affected ocean circulation and/or water column
stratification, but in order to sustain marine productivity for
104 – 105 years and disrupt the global carbon cycle wesuspect that elevated rates of hydrothermal activity indirectly contributed to these events.
5.4. Ocean Circulation, Rising Sea Level,
and Chalk Deposition[52] Reduced export production after the early Albian wasclosely related to rising sea level, global warming, andincreased upper water column stratification. Planktic andbenthic foraminiferal isotopic data from the western North
Atlantic (Blake Nose) show that sea surface temperatures(SSTs) rose rapidly during the Albian (Figure 10). In
addition, isotopic gradients of the upper water column(based on multispecies planktics) as well as planktic-benthic
comparisons show that the water column became increasingly stratified during the Albian. Because temperature hassuch an important effect on metabolic rates [e.g., Brasier,
1995; Vermeij, 1995], increased SSTs would be expected toresult in decreased export production because greater nutrition (food supply) is required by warm water organisms
compared with cooler water taxa. Coupled atmosphereocean general circulation models with two different biogeochemical schemes predict that rising levels of CO2 andglobal warming will cause increased ocean stratification
leading to reduced nutrient supply to the euphotic zone andincreased light efficiency (longer growing season), withboth effects causing a reduction in export production [Boppet al., 2001].
[53] There is a striking parallel between the rapid radiation of angiosperms on land and the diversification of
calcareous plankton during Albian time [e.g., Hickey andDoyle, 1977; Retallack and Dilcher, 1986; Lidgard andCrane, 1988; Crane et al., 1995]. According to Tappan[1986], the evolution of land plants during the Phanerozoic
limited the delivery of land-derived nutrients to the marine
plankton. The partitioning of nutrients between the land andthe marine environment may have been further exacerbatedduring times of higher eustatic sea level owing to thesequestering of organic matter and nutrients in coastal
wetlands and estuaries. In addition, the growth of warm,
oxygen-poor epicontinental seas may have increased therate of denitrification, thereby further limiting the availability of fixed nitrogen for oceanic phytoplankton [Codispoti,
1997; Falkowski, 1997]. Nutrient limitation would havefavored primary productivity by cyanobacteria, coccolithophorids, and other picoautotrophs and nanoflagellates [Laybourn-Parry, 1992; Jumars, 1993; Falkowski et al., 1998].

LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

The initiation of widespread chalk deposition by the late
Albian signals the spread of warm, oligotrophic, stratified
seas, which persisted through Late Cretaceous time. This
marks a fundamental shift in the ecosystem structure ofoceanic plankton communities as near-surface waters of theeuphotic zone became stratified and nutrient-starved.
[54] A widespread upper to middle bathyal (500 – 1000 m)
oxygenation event in the late early or early middle Albianage sediments at Site 1049 in the western North Atlantic andat Site 511 on Falkland Plateau in the South Atlantic
abruptly terminates dysoxic to anoxic conditions at these
two widely separated sites. This event suggests that anintermediate water connection between the North and South
Atlantic Ocean basins may have existed by early Albian
time. If this connection involved the introduction of dense(saline) northern South Atlantic water into the North Atlantic basin at this time [e.g., Arthur and Natland, 1979], theCoriolis effect would have directed the strongest flow ofintermediate or deep water northward along the margin ofnorthwest Africa. The prominent unconformity in the
Albian section of Site 545 may be the consequence oferosion by eastern boundary currents (Figure 2). Later deep
water oxygenation events occurred at or near the end of the
Cenomanian stage [Tucholke and Vogt, 1979; Summerhayes, 1981].
[55] The beginning of the end of black shale deposition inboth the North and South Atlantic Ocean basins was not
simply the consequence of changing trophic dynamics inthe plankton and reduced export production; the end ofdysoxic and anoxic conditions was also related to changing
tectonic gateways and sill depths with the opening of theAtlantic [e.g., de Graciansky et al., 1982; Summerhayes,
1987; Weissert and Lini, 1991]. The propensity for blackshale deposition in the basins of the Atlantic and European
Tethys became progressively minimized during AlbianCenomanian time. The OAEs of the latest Albian andCenomanian/Turonian boundary were triggered by different
tectonic and ocean-climate circumstances than the early
Aptian and latest Aptian– early Albian events. The consequences of long-term rising sea level and global warming
during Albian-Turonian time included (1) creation of epicontinental seas and other gateways to the high latitudes, (2)partitioning of nutrients between flooded coastal plains,epicontinental seas, and the open ocean, (3) changes inoceanic circulation and water mass sources as epicontinental
seas linked Tethys with the higher latitudes and as the
deepwater gateway between the North and South Atlantic
continued to open, (4) changes in the hydrologic cycle andcontinental weathering, and (5) increasing water column
stratification and nutrient partitioning between surface anddeep waters. Taken together, the cumulative effects ofclimate, ocean fertility, ocean circulation, water column
stratification, and carbonate chemistry controlled plankton
populations and their capacity to produce organic carbon
and/or carbonate.

6. Conclusions[56] Turnover (extinction plus speciation) in the planktic
foraminifera closely tracks that of the radiolaria, and both

13 - 23groups of heterotrophic protists display the greatest rates ofturnover at or near the major OAEs. The autotrophic
calcareous nannoplankton were most strongly affected by
the early Aptian OAE1a (Selli event) and the Cenomanian/Turonian boundary OAE2 (Bonarelli event). These two
events were the most widespread of the OAEs; both wereshort-lived and associated with increased marine productivity and active submarine volcanism. The eruption of theOntong Java superplume was a precursor to OAE1a, and the
combined influences of faster seafloor spreading rates and/or other submarine volcanism occurred during the time of
OAE2. We propose that increased marine productivity
during latest Aptian-early Albian OAE1b was also linked
to increased ridge crest volcanism and hydrothermal activity
as indicated by reduced (less radiogenic) 87Sr/86Sr ratios.
However, OAE1b encompasses a protracted interval oforganic carbon burial in contrast to the black shales of
OAE1a and OAE2. A number of different mechanisms
created the conditions necessary to accumulate the multiple
but geographically restricted black shale events of OAE1b.
However, the large carbon isotopic excursions (up to 3%)
during late Aptian-early Albian time (Ticinella bejaouaensis-Hedbergella planispira biozones) indicate that this interval was associated with significant disruptions in the global
carbon cycle.
[57] The OAEs were extraordinary events driven by
extraordinary forcing factors. Each was associated withpronounced carbon isotopic excursions (>1.5 – 2%) andelevated rates of plankton turnover indicating a broad
impact on the ocean-climate system, despite the variable
geographic distribution of black shale deposition, but how
could the ocean be affected on a global scale, and how
could elevated marine productivity be sustained for 104 –105 years? We conclude that the mid-Cretaceous OAEs
were linked to submarine volcanic activity. We hypothesizethat the emplacement of oceanic plateaus and increased
ridge crest hydrothermal activity indirectly helped to fueland sustain the elevated levels of marine productivity during
the OAEs by way of iron fertilization of the water column.
Global warming associated with increased submarine volcanism also contributed to higher productivity by way ofintensified chemical weathering on the land and greater fluxof nutrients to the ocean. Perhaps only OAE1a was directly
triggered by submarine volcanism, but we suggest that
active hydrothermal activity, in particular, indirectly helped
to sustain marine productivity on a global scale during the
OAEs. Transgression also facilitated the production ofmarine organic matter and its burial and preservation ascondensed intervals of black shale deposition.
[58] Changing nutrient availability and/or upper watercolumn structure were the causes of plankton turnover ator near the OAEs. However, excess CO2 associated withsubmarine volcanism may have reduced carbonate availability and seawater pH, thereby contributing to the loss of
calcareous plankton during the early Aptian and latest
Aptian-early Albian OAEs. For example, the heavily calcified nannofossils (the nannoconids) were seriously affectedby the ocean-climate changes associated with OAE1a, andboth nannoconids and the largest and most heavily calcified
planktic foraminifera were either eradicated or severely

13 - 24LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

depleted at the onset of OAE1b. In contrast, the deepestdwelling planktic foraminifera not necessarily the mostheavily calcified taxa were eradicated during OAE2 at the
Cenomanian/Turonian boundary. This event was likely
caused by an expanded oxygen minimum zone associatedwith elevated productivity and/or by the decay of thethermocline due an abrupt deep-sea warming event. The
turnover event associated with OAE1d in the latest Albian
was likewise triggered by a collapse of upper water column
stratification due to surface water cooling or an intermediatewater warming event.
[59] Greater ocean crust production in the late Aptian
initiated a long-term trend of rising sea level and warming
global climate, both of which peaked in the early Turonian(92 –93 Ma). The rising 87Sr/86Sr ratios during the Albianrecord increasing continental weathering rates with rising
sea level and global warming despite the increased ridge
crest volcanism that sustained high global sea level throughmuch of the Late Cretaceous. In addition, the deepening
gateway between the basins of North and South Atlantic
figured prominently in surface and intermediate water ocean
circulation by early Albian time and deep water ventilation
by Cenomanian/Turonian boundary time.
[60] OAE1b in the latest Aptian-early Albian was awatershed event in the evolution of planktic foraminifera.
This interval also marks the beginning of the end of widespread black shale deposition and the initiation of chalk
deposition. According to the hypothesis of Stanley andHardie [1998], greater rates of hydrothermal activity
through the spreading centers altered ocean carbonate
chemistry to favor calcite-secreting plankton and otherorganisms. The marked increases in planktic foraminiferal
size and degree of calcification during the Albian support
this hypothesis. The spread of pelagic carbonate deposition
with rising sea level in the Albian also signals a changing
planktic trophic regime dictated by increased thermal andfertility gradients. Fundamental changes in circulation,water column structure, nutrient partitioning, and plankton
ecosystem dynamics in the post-Aptian ocean favored
stratified, oligotrophic to mesotrophic, diverse, carbonatebased pelagic and benthic ecosystems that came to characterize the widespread chalk seas of the Late Cretaceous.
[61] Acknowledgments. R.M.L. would like to acknowledge thedonors to the American Chemical Society-Petroleum Research Fund forpartial research support. This research used samples provided by the Ocean
Drilling Program (ODP). The ODP is sponsored by the U.S. National
Science Foundation (NSF) and participating countries under managementof Joint Oceanographic Institutions (JOI), Inc. We warmly acknowledge
discussions with Michael Arthur, Steve Burns, Emily CoBabe, Peter Crane,
Rob DeConto, Steve D’Hondt, Elisabetta Erba, Jochen Erbacher, Brian
Huber, Roger Larson, Mitch Malone, Richard Norris, Isabella Premoli
Silva, Paul Wilson, and Richard Yuretich. The paper benefited from thethoughtful reviews of Brian Huber, Steve Nathan, Richard Norris, and Paul
Wilson. A special thanks to Fatima Abrantes and Michael Arthur for theopportunity to present an early version of this paper at the Sixth International Paleoceanographic Conference held in Lisbon in 1998. This paper isdedicated to the memory of William V. Sliter, an international leader in
Cretaceous micropaleontology and our mentor, colleague, and friend.

References
Arthur, M. A., and J. H. Natland, Carbonaceous
sediments in the North and South Atlantic: The
role of salinity in stable stratification of Early
Cretaceous basins, in Deep Drilling Results in
the Atlantic Ocean: Continental Margins andPaleoenvironment, Maurice Ewing Ser., vol. 3,edited by M. Talwani et al., pp. 375 – 401,
AGU, Washington, D. C., 1979.
Arthur, M. A., and I. Premoli Silva, Development of widespread organic carbon-rich strata
in the Mediterranean Tethys, in Nature of Cretaceous Carbon-Rich Facies, edited by S. O.
Schlanger and M. B. Cita, pp. 7 – 54, Academic, San Diego, Calif., 1982.
Arthur, M. A., and B. B. Sageman, Marine blackshales: Depositional mechanisms and environments of ancient deposits, Annu. Rev. Earth
Planet. Sci., 22, 499 – 551, 1994.
Arthur, M. A., and S. O. Schlanger, Cretaceous‘‘oceanic anoxic events’’ as causal factors indevelopment of reef-reservoired giant oilfields, AAPG Bull., 63, 870 – 885, 1979.
Arthur, M. A., W. E. Dean, D. J. Bottjer, and
P. A. Scholle, Rhythmic bedding in Mesozoic-Cenozoic pelagic carbonate sequences:
The primary and diagenetic origin of Milankovitch-like cycles, in Milankovitch and Climate,edited by A. Berger, pp. 191 – 222, D. Riedel,
Norwell, Mass., 1984.
Arthur, M. A., W. E. Dean, and S. O. Schlanger,
Variations in the global carbon cycle during
the Cretaceous related to climate, volcanism,and changes in atmospheric CO2, in The Carbon Cycle and Atmospheric CO2: Natural Variations Archean to Present, Geophys. Monogr.
Ser., vol. 32, edited by E. T. Sundquist and W.
S. Broecker, pp. 504 – 529, AGU, Washington,
D. C., 1985.

Arthur, M. A., S. O. Schlanger, and H. C. Jenkyns, The Cenomanian-Turonian oceanic anoxic event II, paleoceanographic controls onorganic matter production and preservation,in Marine Petroleum Source Rocks, edited by
J. Brooks and A. Fleet, pp. 399 – 418, Geol.
Soc. Spec. Publ., 24, 1987.
Arthur, M. A., W. E. Dean, and L. M. Pratt,Geochemical and climatic effects of increased
marine organic carbon burial at the Cenomanian/Turonian boundary, Nature, 335, 714 –
717, 1988.
Arthur, M. A., H.-J. Brumsack, H. C. Jenkyns,
and S. O. Schlanger, Stratigraphy, geochemistry, and paleoceanography of organic carbonrich Cretaceous sequences, in Cretaceous Resources, Events, and Rhythms, edited by R. N.
Ginsburg and B. Beaudoin, pp. 75 – 119,
Kluwer Acad., Norwell, Mass., 1990.
Azam, F., Microbial control of oceanic carbon
flux: The plot thickens, Science, 280, 694 –
695, 1998.
Azam, F., and R. A. Long, Sea snow microcosms, Nature, 414, 495 – 498, 2001.
Azam, F., T. Fenchel, J. G. Field, J. S. Gray, R. A.
Meyer-Reil, and F. Thingstad, The ecologicalrole of water column microbes in the sea, Mar.
Ecol. Prog. Ser., 10, 257 – 263, 1983.
Baker, E. T., C. R. German, and H. Elderfield,
Hydrothermal plumes over spreading center
axes: Global distributions and geophysical inferences, in Seafloor Hydrothermal Systems:
Physical, Chemical, Biological, and Geological Interactions, Geophys. Monogr. Ser., vol.
91, edited by S. Humphries et al., pp. 47 – 71,
AGU, Washington, D. C., 1995.
Banner, F. T., and D. Desai, A review and revision of the Jurassic-Early Cretaceous Globiger-

inina with special reference to the Aptian
assemblages of Speeton (North Yorkshire, England), J. Micropaleontol., 7, 143 – 185, 1988.
Barron, E. J., A warm, equable Cretaceous: The
nature of the problem, Earth Sci. Rev., 19,
305 – 338, 1983.
Barron, E. J., Global Cretaceous paleogeography—International Geologic Correlation Project 191, Palaeogeogr. Palaeoclimatol.
Palaeoecol., 59, 207 – 216, 1987.
Barron, E. J., and W. H. Peterson, Mid-Cretaceous
ocean circulation: Results from model sensitivity studies, Paleoceanography, 5, 319 – 337,
1990.
Barron, E. J., and W. M. Washington, Warm
Cretaceous climates: High atmospheric CO2
as a plausible mechanism, in The Carbon Cycle and Atmospheric CO2: Natural Variations
Archean to Present, Geophys. Monogr. Ser.,vol. 32, edited by E. T. Sundquist and W. S.
Broecker, pp. 546 – 553, AGU, Washington,
D. C., 1985.
Barron, E. J., P. J. Fawcett, W. H. Peterson,
D. Pollard, and S. L. Thompson, A ‘‘simulation’’ of mid-Cretaceous climate, Paleoceanography, 10, 953 – 962, 1995.
Barron, E. J., W. W. Hay, and S. Thompson, The
hydrologic cycle: A major variable during
Earth history, Palaeogeogr. Palaeoclimatol.
Palaeoecol., 75, 157 – 174, 1989.
Barron, E. J., W. H. Peterson, S. L. Thompson,and D. Pollard, Past climate and the role ofocean heat transport: Model simulations for
the Cretaceous, Paleoceanography, 8, 785 –
798, 1993.
Behrenfeld, M. J., and Z. S. Kolber, Widespreadiron limitation of phytoplankton in the South
Pacific Ocean, Science, 283, 840 – 843, 1999.

LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION
Berger, W. H., and L. Diester-Haas, Paleoproductivity: The benthic/planktonic ratio in foraminifera as a productivity index, Mar. Geol.,81, 15 – 25, 1988.
Berger, W. H., J. S. Killingley, and E. Vincent,Stable isotopes in deep-sea carbonates: Box
core ERDC-92, west equatorial Pacific, Oceanol. Acta, 1, 203 – 216, 1978.
Bopp, L., P. Monfray, O. Aumont, J.-L. Dufresne, H. Le Treut, G. Madec, L. Terray, andJ. C. Orr, Potential impact of climate change
on marine export production, Global Biogeochem. Cycles, 15, 81 – 99, 2001.
BouDagher-Fadel, M. K., F. T. Banner, and J. E.
Whittaker, The Early Evolutionary History of
Planktonic Foraminifera, 269 pp., Chapman
and Hall, New York, 1997.
Bralower, T. J., Calcareous nannofossil biostratigraphy and assemblages of the CenomanianTuronian boundary interval: Implications forthe origin and timing of oceanic anoxia, Paleoceanography, 3, 275 – 316, 1988.
Bralower, T. J., and J. A. Bergen, CenomanianSantonian calcareous nannofossil biostratigraphy of a transect of cores drilled across theWestern Interior Seaway, in Stratigraphy andPaleoenvironments of the Cretaceous WesternInterior Seaway, USA, Concepts in Sedimentol.
Paleontol., vol. 6, edited by W. E. Dean and
M. A. Arthur, pp. 59 – 77, Soc. Sediment.
Geol., Tulsa, Okla., 1998.
Bralower, T. J., and W. G. Siesser, Cretaceous
calcareous nannofossil stratigraphy of ODP
Sites 761, 762, and 763, Exmouth and Wombat
Plateaus, N.W, Australia, Proc. Ocean Drill.
Program Sci. Results, 122, 529 – 566, 1992.
Bralower, T. J., W. V. Sliter, M. A. Arthur, R. M.
Leckie, D. Allard, and S. O. Schlanger, Dysoxic/anoxic episodes in the Aptian-Albian(Early Cretaceous), in The Mesozoic Pacific:
Geology, Tectonics and Volcanism, Geophys.
Monogr. Ser., vol. 77, edited by M. S. Pringle
et al., pp. 5 – 37, AGU, Washington, D. C.,
1993.
Bralower, T. J., M. A. Arthur, R. M. Leckie, W. V.
Sliter, D. Allard, and S. O. Schlanger, Timing
and paleoceanography of oceanic dysoxia/anoxia in the late Barremian to early Aptian, Palaios, 9, 335 – 369, 1994.
Bralower, T. J., R. M. Leckie, W. V. Sliter, and
H. R. Thierstein, An integrated Cretaceous microfossil biostratigraphy, in Geochronology,
Time Scales, and Global Stratigraphic Correlation, edited by W. A. Berggren et al., 65 – 79,
Spec. Publ. SEPM Soc. Sedment. Geol., 54,
1995.
Bralower, T. J., P. D. Fullagar, C. K. Paull, G. S.
Dwyer, and R. M. Leckie, Mid-Cretaceousstrontium-isotope stratigraphy of deep-sea sections, Geol. Soc. Am. Bull., 109, 1421 – 1442,
1997.
Bralower, T. J., E. CoBabe, B. Clement, W. V.
Sliter, C. L. Osburn, and J. Longoria, The record of global change in mid-Cretaceous (Barremian-Albian) sections from the Sierra
Madre, northeastern Mexico, J. ForaminiferalRes., 29, 418 – 437, 1999.
Brasier, M. D., Fossil indicators of nutrient levels, 1, Eutrophication and climate change, inMarine Palaeoenvironmental Analysis from
Fossils, edited by D. W. Bosence and P. A.
Allison, London, Geol. Soc. Spec. Publ.,83,
113 – 132, 1995.
Brass, G. W., J. P. Southam, and W. H. Peterson,
Warm saline bottom water in the ancient
ocean, Nature, 269, 620 – 623, 1982.
Bréhéret, J.-G., Glauconitization episodes inmarginal settings as echoes of mid-Cretaceous

anoxic events in the Vocontian basin (SEFrance), Modern and Ancient Continental
Shelf Anoxia, edited by R. V. Tyson and T. H.
Peterson, Geol. Soc. Spec. Publ.,58, 415 – 425,
1991.
Bréhéret, J.-G., The mid-Cretaceous organic-richsediments from the Vocontian zone of theFrench Southeast Basin, in Hydrocarbon andPetroleum Geology of France, edited by A.
Mascle, pp. 295 – 320, Springer-Verlag, New
York, 1994.
Bréhéret, J.-G., and M. Delamette, Faunal fluctuations related to oceanographical changes inthe Vocontian basin (SE France) during Aptian-Albian time, Geobios Mem. Spec., 11,267 – 277, 1989.
Bréhéret, J.-G., M. Caron, and M. Delamette,
Niveaux riches en matière organique dansl’Albien vocontien; quelques caractères du paléoenvironnement essai d’interprétation génétique, in Les Couches Riches en MatièreOrganique et leurs Conditions de Dépôt, edited by J.-G. Bréhéret, Documents B.R.G.M.,
110, 141 – 191, 1986.
Bujak, J. P., and G. L. Williams, Dinoflagellate
diversity through time, Mar. Micropaleontol.,
4, 1 – 12, 1979.
Capone, D. G., J. P. Zehr, H. W. Paerl, B. Bergman, and E. J. Carpenter, Trichodesmium, a
globally significant marine cyanobacterium,
Science, 276, 1221 – 1229, 1997.
Caron, M., Cretaceous planktonic foraminifera,in Plankton Stratigraphy, edited by H. M. Bolli et al., pp. 17 – 86, Cambridge Univ. Press,
New York, 1985.
Caron, M., and P. Homewood, Evolution of earlyplanktonic foraminifers, Mar. Micropaleontol.,
7, 453 – 462, 1983.
Caron, M., F. Robaszynski, F. Amedro, F. Baudin, J.-F. Deconinck, P. Hochuli, K. SalisPerch Nielsen, and N. Tribovillard, Estimationde la durée de l’événement anoxique global au
passage Cénomanien/Turonien. Approche cyclostratigraphique dans la formation Bahloul
en Tunisie centrale, Bull. Soc. Geol. Fr., 170,
145 – 160, 1999.
Chamberlin, T. C., On a possible reversal of deepsea circulation and its influence on geologic
climate, J. Geol., 14, 363 – 373, 1906.
Cho, B. C., and F. Azam, Major role of bacteria
in biogeochemical fluxes in the ocean’s interior, Nature, 332, 441 – 443, 1988.
Clarke, L. J., and H. C. Jenkyns, New oxygenisotope evidence for long-term Cretaceous climatic change in the Southern Hemisphere,
Geology, 27, 699 – 702, 1999.
Coale, K. H., S. E. Fitzwatere, R. M. Gordon,
K. S. Johnson, and R. T. Barber, Control ofcommunity growth and export production by
upwelled iron in the equatorial Pacific Ocean,
Nature, 379, 621 – 624, 1996a.
Coale, K. H., et al., A massive phytoplankton
bloom induced by an ecosystem-scale iron fertilization experiment in the equatorial Pacific
Ocean, Nature, 383, 495 – 501, 1996b.
Coccioni, R., and S. Galeotti, Orbitally induced
cycles in benthonic foraminiferal morphogroups and trophic structure distribution patterns from the late Albian ‘‘Amadeus Segment’’(central Italy), J. Micropaleontol., 12, 227 –
239, 1993.
Coccioni, R., and I. Premoli Silva, Planktonic
foraminifera from the Lower Cretaceous of
Rio Argos sections (southern Spain) and biostratigraphic implications, Cretaceous Res., 15,645 – 687, 1994.
Coccioni, R., E. Erba, and I. Premoli Silva, Barremian-Aptian calcareous plankton biostrati-

13 - 25

graphy from Gorgo Cerbara section (Marche,
central Italy) and implications for plankton
evolution, Cretaceous Res., 13, 517 – 537,
1992.
Codispoti, L. A., Phosphorous vs. nitrogen limitation of new and export production, in Productivity of the Ocean: Present and Past,
edited by W. H. Berger et al., pp. 377 – 394,
John Wiley, 1989.
Codispoti, L. A., The limits to growth, Nature,
387, 237 – 238, 1997.
Coffin, M. F., and O. Eldholm, Large igneous
provinces: Crustal structure, dimensions, andexternal consequences, Rev. Geophys., 32,
1 – 36, 1994.
Cool, T. E., Sedimentological evidence concerning the paleoceanography of the Cretaceous
western North Atlantic Ocean, Palaeogeogr.
Palaeoclimatol. Palaeoecol., 39, 1 – 35, 1982.
Corfield, R. M., M. A. Hall, and M. D. Brasier,Stable isotope evidence for foraminiferal habitats during the development of the Cenomanian-Turonian oceanic anoxic event, Geology,
18, 175 – 178, 1990.
Cotillon, P., and M. Rio, Cyclic sedimentation in
the Cretaceous of Deep Sea Drilling Project
Sites 535 and 540 (Gulf of Mexico), 534 (Central Atlantic), and in the Vocontian Basin(France), Initial Rep. Deep Sea Drill. Proj.,77, 334 – 377, 1984.
Crane, P. R., E. M. Friis, and K. R. Pedersen, The
origin and early diversification of angiosperms, Nature, 374, 27 – 33, 1995.
Dean, W. E., J. V. Gardner, L. F. Jansa, P. Cepek,
and D. Seibold, Cyclic sedimentation along thecontinental margin of northwest Africa, Initial
Rep. Deep Sea Drill. Proj., 41, 965 – 989,
1978.
de Boer, P. L., Cyclicity and storage of organic
matter in middle Cretaceous pelagic sediments,
in Cyclic and Event Stratification, edited by G.
Einsele and A. Seilacher, pp. 456 – 475,
Springer-Verlag, New York, 1982.
de Boer, P. L., and A. A. H. Wonders, Astronomically induced thythmic bedding, in Milankovitch and Climate, part 1, edited by A.
L. Berger et al., pp. 177 – 190, D. Riedel, Norwell, Mass., 1984.
de Graciansky, P. C., et al., Les formations d’âgeCrétacé de l’Atlantique Nord et leur metièreorganique: Paleogéographie et milieux de depot, Rev. Inst. Fr. Petrole, 37, 275 – 337, 1982.
D’Hondt, S., and M. A. Arthur, Interspecific variation in stable isotope signals of Maastrichtian
planktonic foraminifera, Paleoceanography,
10, 123 – 135, 1995.
Deuser, W. G., E. H. Ross, C. Hemleben, and M.
Spindler, Seasonal changes in species composition, number, mass, size, and isotopic composition of planktonic foraminifera settling
into the deep Sargasso Sea, Palaeogeogr. Palaeoclimatol. Palaecol., 33, 103 – 127, 1981.
Eicher, D. L., Cenomanian and Turonian planktonic foraminifera from the western interior of
the United States, in Proceedings of the FirstInternational Conference on Planktonic Microfossils, vol. 2, edited by P. Brönnimannand H. H. Renz, pp. 163 – 174, E. J. Brill, Cologne, Germany, 1969.
Erba, E., Calcareous nannofossil distribution inpelagic rhythmic sediments (Aptian-Albian
Piobbico core, central Italy), Riv. Ital. Paleontol. Strat., 97, 455 – 484, 1992.
Erba, E., Nannofossils and superplumes: The
early Aptian ‘‘nannoconid crisis’’, Paleoceanography, 9, 483 – 501, 1994.
Erba, E., I. Premoli Silva, and D. K. Watkins,
Cretaceous calcareous plankton stratigraphy

13 - 26

LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

of Sites 872 through 879, Proc. Ocean Drill.
Program Sci. Results, 144, 157 – 169, 1996.
Erba, E., J. E. T. Channell, M. Claps, C. Jones,
R. Larson, B. Opdyke, I. Premoli Silva, A.
Riva, G. Salvini, and S. Torricelli, Integrated
stratigraphy of the Cismon Apticore (southern
Alps, Italy): A ‘‘reference section’’ for the Barremian-Aptian interval at low latitudes, J. Foramaminiferal Res., 29, 371 – 391, 1999.
Erbacher, J., and J. Thurow, Influence of oceanic
anoxic events on the evolution of mid-Cretaceous radiolaria in the North Atlantic and western Tethys, Mar. Micropaleontol., 30, 139 –
158, 1997.
Erbacher, J., J. Thurow, and R. Littke, Evolution
patterns of radiolaria and organic matter variations: A new approach to identify sea levelchanges in mid-Cretaceous pelagic environments, Geology, 24, 499 – 502, 1996.
Erbacher, J., W. Gerth, G. Schmiedl, and C.
Hemleben, Benthic foraminiferal assemblagesof late Aptian-early Albian black shale intervals in the Vocontian Basin, SE France, Cretaceous Res., 19, 805 – 826, 1998.
Erbacher, J., C. Hemleben, B. T. Huber, and M.
Markey, Correlating environmental changes
during early Albian oceanic anoxic event 1B
using benthic foraminiferal paleoecology, Mar.
Micropaleontol., 38, 7 – 28, 1999.
Erbacher, J., B. T. Huber, R. D. Norris, and M.
Markey, Increased thermohaline stratification
as a possible cause for an ocean anoxic eventin the Cretaceous period, Nature, 409, 325 –
327, 2001.
Erez, J., and B. Luz, Experimental paleotemperature equation for planktonic foraminifera, Geochim. Cosmochim. Acta, 47, 1025 – 1031,
1983.
Fairbanks, R. G., and P. H. Wiebe, Foraminifera
and chlorophyll maximum: Vertical distribution, seasonal succession, and paleoceanographic significance, Science, 209, 1524 –
1526, 1980.
Fairbanks, R. G., M. Sverdlove, R. Free, P. H.
Wiebe, and A. W. Bé, Vertical distribution andisotopic fractionation of living planktonic foraminifera from the Panama Basin, Nature, 298,
841 – 844, 1982.
Falkowski, P. G., Evolution of the nitrogen cycle
and its influence on the biological sequestration of CO2 in the ocean, Nature, 387, 272 –
275, 1997.
Falkowski, P. G., D. Ziemann, Z. Kolber, and P. K.
Bienfang, Role of eddy pumping in enhancing
primary productivity in the ocean, Nature, 353,
55 – 58, 1991.
Falkowski, P. G., R. T. Barber, and V. Smetacek,
Biochemical controls and feedbacks on ocean
primary productivity, Science, 281, 200 – 206,
1998.
Fassell, M. L., and T. J. Bralower, Warm, equable mid-Cretaceous: Stable isotope evidence,in Evolution of the Cretaceous Ocean-Climate
System, edited by E. Barrera and C. C. Johnson, Spec. Pap. Geol. Soc. Am., 332, 121 –
142, 1999.
Fenchel, T., Marine bugs and carbon flow,
Science, 292, 2444 – 2445, 2001.
Ferguson, K. M., R. T. Gregory, and A. Constantine, Lower Cretaceous (Aptian-Albian) secular changes in the oxygen and carbon isotope
record from high paleolatitude, fluvial sediments, southeast Australia: Comparisons to
the marine record, in Evolution of the Cretaceous Ocean-Climate System, edited by E. Barrera and C. C. Johnson, Spec. Pap. Geol. Soc.
Am., 332, 59 – 72, 1999.
Fischer, A. G., Climate rhythms recorded in stra-ta, Annu. Rev. Earth Planet. Sci, 14, 351 – 376,
1986.
Fischer, A. G., and M. A. Arthur, Secular variations in the pelagic realm, in Deepwater Carbonate Environments, edited by H. E. Cookand P. Enos, Spec. Publ. Soc. Econ. Paleontol.
Mineral., 25, 19 – 50, 1977.
Föllmi, K. B., H. Weissert, M. Bisping, and H.
Funk, Phosphogenesis, carbon-isotope stratigraphy, and carbonate platform evolution
along the Lower Cretaceous northern Tethyan
margin, AAPG Bull., 106, 729 – 746, 1994.
Frakes, L. A., Estimating the global thermal statefrom Cretaceous sea surface and continental
temperature data, in Evolution of the Cretaceous Ocean-Climate System, edited by E. Barrera and C. C. Johnson, Spec. Pap. Geol. Soc.
Am., 332, 49 – 57, 1999.
Frakes, L. A., and J. E. Francis, A guide to Phanerozoic cold polar climates from high-latitude
ice-rafting in the Cretaceous, Nature, 333,
547 – 549, 1988.
Frakes, L. A., J. E. Francis, and J. I. Syktus,
Climate Modes of the Phanerozoic, 274 pp.,
Cambridge Univ. Press, New York, 1992.
Frey, F., M. Coffin, and P. J. Wallace, Origin andevolution of a submarine large igneous province: The Kerguelen Plateau and Broken
Ridge, southern Indian Ocean (abstract), Eos
Trans AGU, 80(46), F1103, Fall Meet. Suppl.,
1999.
Froelich, P. N., M. L. Bender, N. A. Luedtke,
G. R. Heath, and T. DeVries, The marine phosphorous cycle, Am. J. Sci., 282, 474 – 511,
1982.
Frost, B. W., Phytoplankton bloom on iron rations, Nature, 383, 475 – 476, 1996.
Frush, M. P., and D. L. Eicher, Cenomanian andTuronian foraminifera and paleoenvironmentsin the Big Bend region of Texas and Mexico,in The Cretaceous System in the Western Interior of North America, edited by W. G. E. Caldwell et al., Geol. Assoc. Can. Spec. Pap., 13,277 – 301, 1975.
Gale, A. S., H. C. Jenkyns, W. J. Kennedy, andR. M. Corfield, Chemostratigraphy versus
biostratigraphy: Data from around the Cenomanian-Turonian boundary, J. Geol. Soc. London, 150, 29 – 32, 1993.
Gale, A. S., A. B. Smith, N. E. A. Monks, J. A.
Young, A. Howard, D. S. Wray, and J. M.
Huggett, Marine biodiversity through the late
Cenomanian-early Turonian: Palaeoceanographic controls and sequence stratigraphic
biases, J. Geol. Soc. London, 157, 745 – 757,
2000.
Gasperi, J. T., and J. P. Kennett, Isotopic evidence for depth stratification and paleoecologyof Miocene planktonic foraminifera: Westernequatorial Pacific DSDP Site 289, in Pacific
Neogene-Environment, Evolution, and Events,
edited by R. Tsuchi and J. C. Ingle, pp. 117 –
147, Univ. of Tokyo Press, Tokyo, 1992.
Gersonde, R., and D. M. Harwood, Lower Cretaceous diatoms from ODP Leg 113 Site 693(Weddell Sea), part 1, Vegetative cells, Proc.
Ocean Drill. Program Sci. Results, 113, 365 –
402, 1990.
Glover, H. E., B. B. Prezelin, L. Campbell, M.
Wyman, and C. Garside, A nitrate-dependent
Synechococcus bloom in surface Sargasso Sea
water, Nature, 331, 161 – 163, 1988.
Goldstein, S. T., Foraminifera: A biological overview, in Modern Foraminifera, edited by B. K.
Sen Gupta, pp. 37 – 55, Kluwer Acad., Norwell, Mass., 1999.
Gradstein, F. M., Biostratigraphy of Lower Cretaceous Blake Nose and Blake-Bahama Basin

foraminifers, DSDP Leg 44, western North
Atlantic Ocean, Initial Rep. Deep Sea Drill.
Proj., 44, 663 – 701, 1978.
Gradstein, F. M., F. P. Agterberg, J. G. Ogg, J.
Hardenbol, P. van Veen, J. Thierry, and Z.
Huang, A Mesozoic time scale, J. Geophys.
Res., 99, 24,051 – 24,074, 1994.
Gröcke, D. R., S. P. Hesselbo, and H. C. Jenkyns, Carbon-isotope composition of lower
Cretaceous fossil wood: Ocean-atmospherechemistry and relation to sea-level change,
Geology, 27, 155 – 158, 1999.
Haig, D. W., and D. A. Lynch, A late early Albian marine transgressive pulse over northeastern Australia, precursor to epeiric basin anoxia:
Foraminiferal evidence, Mar. Micropaleontol.,
22, 311 – 362, 1993.
Hallam, A., and M. J. Bradshaw, Bituminous
shales and oolitic ironstones as indicators oftransgressions and regressions, J. Geol. Soc.
London, 136, 157 – 164, 1979.
Haq, B. U., Transgressions, climatic change, andthe diversity of calcareous nannoplankton,
Mar. Geol., 15, 25 – 30, 1973.
Haq, B. U., J. Hardenbol, and P. R. Vail, Mesozoic and Cenozoic chronostratigraphy and cycles of sea-level change, in Sea-Level Changes:
An Integrated Approach, edited by C. K. Wilgus et al., Spec. Publ. Soc. Econ. Paleontol.
Mineral., 42, 71 – 108, 1988.
Hardie, L. A., Secular variation in seawaterchemistry: An explanation for the coupled secular variation in the mineralogies of marine
limestones and potash evaporites over the past
600 m.y., Geology, 24, 279 – 283, 1996.
Hart, M. B., A water depth model for the evolution of the planktonic Foraminiferida, Nature,
286, 252 – 254, 1980.
Hart, M. B., and K. C. Ball, Late Cretaceous
anoxic events, sea-level changes, and the evolution of the planktonic foraminifera, in North
Atlantic Paleoceanography, edited by C. P.
Summerhayes and N. J. Shackleton, Geol.
Soc Spec. Publ., 21, 67 – 78, 1986.
Harwood, D. M., and R. Gersonde, Lower Cretaceous diatoms from ODP Leg 113 Site 693(Weddell Sea), part 2, Resting spores, Chrysophycean cysts, and endoskeletal dinoflagellates, and notes on the origin of diatoms,
Proc. Ocean Drill. Program Sci. Results,
113, 403 – 426, 1990.
Harwood, D. M., and V. A. Nikolaev, Cretaceous
diatoms: Morphology, taxonomy, biostratigraphy, in Siliceous Microfossils, Short Courses
in Paleontol., vol. 8, edited by C. D. Blome etal., pp. 81 – 106, Paleontol. Soc., Knoxville,
Tenn., 1995.
Hay, W. W., Cretaceous paleoceanography, Geol.
Carpathica, 46, 257 – 266, 1995.
Hay, W. W., and R. M. DeConto, Comparison ofmodern and Late Cretaceous meridional energy transport and oceanology, in Evolutionof the Cretaceous Ocean-Climate System, edited by E. Barrera and C. C. Johnson, Spec.
Pap. Geol. Soc. Am., 332, 283 – 300, 1999.
Hays, J. D., and W. C. Pitman III, Lithosphericplate motion, sea level changes and climaticand ecological consequences, Nature, 246,
18 – 22, 1973.
Hemleben, C., M. Spindler, and O. R. Anderson,
Modern Planktonic Foraminifera, SpringerVerlag, New York, 1989.
Herbert, T. D., and A. G. Fischer, Milankovitch
climate origin of mid-Cretaceous black shalerhythms, central Italy, Nature, 321, 739 – 743,
1986.
Herbert, T. D., R. F. Stallard, and A. G. Fischer,Anoxic events, productivity rhythms, and theLECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION
orbital signature in a mid-Cretaceous deep sea
sequence from central Italy, Paleoceanography, 1, 495 – 506, 1986.
Herguera, J. C., and W. H. Berger, Paleoproductivity from benthic foraminifera abundance:
Glacial to postglacial change in the west-equatorial Pacific, Geology, 19, 1173 – 1176, 1991.
Hickey, L. J., and J. A. Doyle, Early Cretaceous
fossil evidence for angiosperm evolution, Bot.
Rev., 43, 3 – 104, 1977.
Hilbrecht, H., H.-W. Hubberten, and H. Öberhänsli, Biogeography of planktonic foraminifera and regional carbon isotope variations:
Productivity and water masses in Late Cretaceous Europe, Palaeogeogr. Palaeoclimatol.
Palaeoecol., 92, 407 – 421, 1992.
Hinz, K., et al., Initial Reports of the Deep Sea
Drilling Project, vol. 79, U.S. Govt. Printing
Off., Washington, D.C., 1984.
Hochuli, P., A. P. Menegatti, H. Weissert, E.
Erba, and I. Premoli Silva, High-productivity
and cooling episodes in the early Aptian Alpine Tethys, Geology, 27, 657 – 660, 1999.
Holbourn, A., and W. Kuhnt, No extinctions during oceanic anoxic event 1b: The Aptian-Albian benthic foraminiferal record of ODP Leg
171, in Western North Atlantic Palaeogene
and Cretaceous Palaeoceanography, edited
by D. Kroon, R. D. Norris, and A. Klaus, Geol.
Soc. Spec. Publ., 183, 73 – 92, 2001.
Huber, B. T., D. A. Hodell, and C. P. Hamilton,
Middle-Late Cretaceous climate of the southern high latitudes: Stable isotopic evidence forminimal equator-to-pole thermal gradients,
Geol. Soc. Am. Bull., 107, 1164 – 1191, 1995.
Huber, B. T., R. M. Leckie, R. D. Norris, T. J.
Bralower, and E. CoBabe, Foraminiferal assemblage and stable isotopic change acrossthe Cenomanian-Turonian boundary in thesubtropical North Atlantic, J. ForaminiferalRes., 29, 392 – 417, 1999.
Huber, B. T., R. D. Norris, and K. G. MacLeod,
Deep-sea paleotemperature record of extreme
warmth during the Cretaceous, Geology, 30,
123 – 126, 2002.
Ingram, B. L., R. Coccioni, A. Montanari, andF. M. Richter, Strontium isotopic compositionof mid-Cretaceous seawater, Science, 264,
546 – 550, 1994.
Jahren, A. H., and N. C. Arens, Methane hydratedissociation implicated in Aptian OAE events,
Geol. Soc. Am. Abstr. Programs, 30, 52, 1998.
Jahren, A. H., N. C. Arens, G. Sarmiento, J.
Guerrero, and R. Amundson, Terrestrial recordof methane hydrate dissociation in the Early
Cretaceous, Geology, 29, 159 – 162, 2001.
Jarvis, I., G. A. Carson, M. K. E. Cooper, M. B.
Hart, P. N. Leary, B. A. Tocher, D. Horne, and
A. Rosenfeld, Microfossil assemblages and the
Cenomanian-Turonian (Late Cretaceous) oceanic anoxic event, Cretaceous Res., 9, 3 – 103,
1988.
Jenkyns, H. C., Cretaceous anoxic events: From
continents to oceans, J. Geol. Soc. London,
137, 171 – 188, 1980.
Jenkyns, H. C., Carbon-isotope stratigraphy andpaleoceanographic significance of the lower
Cretaceous shallow-water carbonates of Resolution Guyot, Mid-Pacific Mountains, Proc.
Ocean Drill. Program Sci. Results, 143, 99 –108, 1995.
Jenkyns, H. C., A. S. Gale, and R. M. Corfield,
Carbon and oxygen-isotope stratigraphy of the
English chalk and Italian scaglia and its paleoclimatic significance, Geol. Mag., 131, 1 – 34,
1994.
Johnson, C. C., E. J. Barron, E. G. Kauffman,
M. A. Arthur, P. J. Fawcett, and M. K. Yasuda,

Middle Cretaceous reef collapse linked toocean heat transport, Geology, 24, 376 – 380,
1996.
Jones, C. E., and H. C. Jenkyns, Seawater strontium isotopes, oceanic anoxic events, and seafloor hydrothermal activity in the Jurassic andCretaceous, Am. J. Sci., 301, 112 – 149, 2001.
Jones, C. E., H. C. Jenkyns, A. L. Coe, and S. P.
Hesselbo, Strontium isotopic variations in Jurassic and Cretaceous seawater, Geochim. Cosmochim. Acta, 58, 3061 – 3074, 1994.
Jumars, P. A., Concepts in Biological Oceanography, An Interdisciplinary Primer, 348 pp.,Oxford Univ. Press, New York, 1993.
Kaiho, K., Phylogeny of deep-sea calcareous trochospiral benthic foraminifera: Evolution anddiversification, Micropaleontology, 44, 291 –
311, 1998.
Kaiho, K., Evolution in the test size of deep-sea
benthic foraminifera during the past 120 m.y.,Mar. Micropaleontol., 37, 53 – 65, 1999.
Kaiho, K., and T. Hasegawa, End-Cenomanian
benthic foraminiferal extinctions and dysoxic
events in the northwestern Pacific Ocean margin, Palaeogeogr. Palaeoclimatol. Palaeoecol., 111, 29 – 43, 1994.
Kaiho, K., O. Fujiwara, and I. Motoyama, MidCretaceous faunal turnover of intermediatewater benthic foraminifera in the northwestern
Pacific Ocean margin, Mar. Micropaleontol.,
23, 13 – 49, 1993.
Kauffman, E. G., Paleobiogeography and evolutionary response dynamic in the CretaceousWestern Interior Seaway of North America,
in Jurassic-Cretaceous Biochronology and Paleogeography of North America, edited by G.
E. G. Westermann, Geol. Assoc. Can. Spec.
Pap., 27, 273 – 306, 1984.
Kaufmann, E. G., and W. G. E. Caldwell, Thewestern interior basin in space and time, in
Evolution of the Western Interior Basin, edited
by W. G. E. Caldwell and E. G. Kauffman,
Geol. Assoc. Can Spec. Pap., 39, 1 – 30, 1993.
Kennett, J. P., and L. D. Stott, Abrupt deep-sea
warming, paleoceanographic changes andbenthic extinction at the end of the Paleocene,
Nature, 353, 225 – 229, 1991.
Kerr, A. C., Oceanic plateau formation: A causeof mass extinction and black shale deposition
around the Cenomanian-Turonian boundary,
J. Geol. Soc. London, 155, 619 – 626, 1998.
Kirchman, D. L., Microbial ferrous wheel, Nature, 383, 303 – 304, 1996.
Kolber, Z. S., F. G. Plumley, A. S. Lang, J. T.
Beatty, R. E. Blankenship, C. L. VanDover, C.
Vetriani, M. Koblizek, C. Rathgeber, and P. G.
Falkowski, Contribution of aerobic photoheterotrophic bacteria to the carbon cycle in theocean, Science, 292, 2492 – 2495, 2001.
Kominz, M. A., Oceanic ridge volumes and sealevel change—An error analysis, in Interregional
Unconformities and Hydrocarbon Accumulation, edited by J. S. Schlee, AAPG Mem., 36,
109 – 127, 1984.
Kuypers, M. M. M., R. D. Pancost, and J. S.
Sinninghe Damsté, A large and abrupt fall in
atmospheric CO2 concentration during Cretaceous times, Nature, 399, 342 – 345, 1999.
Kuypers, M. M. M., P. Blokker, J. Erbacher, H.
Kinkel, R. D. Pancost, S. Schouten, and J. S.
Sinninghe Damsté, Massive expansion of marine Archaea during a mid-Cretaceous oceanic
anoxic event, Science, 293, 92 – 94, 2001.
Larson, R. L., Geological consequences of superplumes, Geology, 19, 963 – 966, 1991a.
Larson, R. L., Latest pulse of Earth: Evidence fora mid-Cretaceous superplume, Geology, 19,547 – 550, 1991b.

13 - 27

Larson, R. L., Superplumes and ridge interactions between Ontong Java and Manihiki Plateaus and the Nova-Canton Trough, Geol., 25,779 – 782, 1997.
Larson, R. L., and E. Erba, Onset of the mid-Cretaceous greenhouse in the Barremian-Aptian:Igneous events and the biological, sedimentary,and geochemical responses, Paleoceanography, 14, 663 – 678, 1999.
Larson, R. L., and C. Kincaid, Onset of midCretaceous volcanism by elevation of the 670
km thermal boundary layer, Geology, 24,
551 – 554, 1996.
Larson, R. L., and W. C. Pitman III, World-wide
correlation of Mesozoic magnetic anomalies,
and its implications, Geol. Soc. Am. Bull., 83,3645 – 3662, 1972.
Laybourn-Parry, J., Protozoan Plankton Ecology,
231 pp., Chapman and Hall, New York, 1992.
Leckie, R. M., Mid-Cretaceous planktonic foraminiferal biostratigraphy off central Morocco,
Deep Sea Drilling Project Leg 79, Sites 545
and 547, Initial Rep. Deep Sea Drilling Proj.,
79, 579 – 620, 1984.
Leckie, R. M., Foraminifera of the CenomanianTuronian boundary interval, Greenhorn Formation, Rock Canyon Anticline, Pueblo, Colorado,in Fine-Grained Deposits and Biofacies of theCretaceous Western Interior Seaway: Evidenceof Cyclic Sedimentary Processes, Field Trip
Guidebk., vol. 4, edited by L. M. Pratt et al.,pp. 139 – 149, Soc. Econ. Paleontol. Mineral.,Tulsa, Okla., 1985.
Leckie, R. M., Paleoecology of mid-Cretaceousplanktonic foraminifera: A comparison of open
ocean and epicontinental sea assemblages, Micropaleontology, 33, 164 – 176, 1987.
Leckie, R. M., An oceanographic model for the
early evolutionary history of planktonic foraminifera, Palaeogeogr. Palaeoclimatol. Palaeoecol., 73, 107 – 138, 1989.
Leckie, R. M., R. F. Yuretich, O. L. O. West, D.
Finkelstein, and M. Schmidt, Paleoceanography of the southwestern Western Interior Sea
during the time of the Cenomanian-Turonianboundary (Late Cretaceous), in Stratigraphy
and Paleoenvironments of the Cretaceous Western Interior Seaway, USA, Concepts in Sedimentol. Paleontol., vol. 6, edited by W. E.
Dean and M. A. Arthur, pp. 101 – 126, Soc.
Sediment. Geol., Tulsa, Okla., 1998.
Lee, J. J., Nutrition and physiology of the foraminifera, in Biochemistry and Physiology ofProtozoa, vol. 3, edited by M. Levandowsky,
and S. H. Hutner, pp. 43 – 66, Academic, San
Diego, Calif., 1980.
Lidgard, S., and P. R. Crane, Quantitative analyses of the early angiosperm radiation, Nature, 331, 344 – 346, 1988.
Lipps, J. H., Plankton evolution, Evolution, 24,
1 – 22, 1970.
Lipps, J. H., Biotic interactions in benthic foraminifera, in Biotic Interactions in Recent andFossil Benthic Communities, edited by M. J. S.
Trevesz and P. L. McCall, pp. 331 – 376, Plenum, New York, 1983.
Longhurst, A., Iron grip on export production,
Nature, 379, 585 – 586, 1996.
Longoria, J. F., Stratigraphic, morphologic, andtaxonomic studies of Aptian planktonic foraminifera, Rev. Esp. Micropaleontol., num. extraordinario, 134 pp., 1974.
Longoria, J. F., and M. A. Gamper, Albian planktonic foraminiferaa from the Sabinas basin ofnorthern Mexico, J. Foraminiferal Res., 7,
196 – 215, 1977.
Loutit, T. S., J. Hardenbol, P. R. Vail, and G. R.
Baum, Condensed sections: The key to age

13 - 28

LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION

determination and correlation of continental
margin sequences, in Sea-Level Changes: An
Integrated Approach, edited by C. K. Wilgus
et al., pp. 183 – 213, Soc. Econ. Mineral. Paleontol., Tulsa, Okla., 1988.
Ludwig, W. J., et al., Initial Reports of the DeepSea Drilling Prog., vol. 71, U.S. Govt. Print.
Off., Washington, D.C., 1983.
Mahoney, J. J., M. Storey, R. A. Duncan, K. J.
Spencer, and M. Pringle, Geochemistry andage of the Ontong Java Plateau, in The Mesozoic Pacific: Geology, Tectonics, and Volcanism, Geophys. Monogr. Ser., vol. 77, edited by
M. Pringle et al., pp. 233 – 261, AGU, Washington, D.C., 1993.
Mahoney, J. J., et al., Initial Reports of the Ocean
Drilling Program, 192, Ocean Drill. Program,College Station, Tex., in press, 2002.
Martin, J. H., Glacial-interglacial CO2 change:
The iron hypothesis, Paleoceanography, 5,
1 – 13, 1990.
Martin, J. H., and S. E. Fitzwater, Iron deficiency
limits phytoplankton growth in the northeast
Pacific subarctic, Nature, 331, 341 – 343,
1988.
Martin, J. H., et al., Testing the iron hypothesis inecosystems of the equatorial Pacific Ocean,
Nature, 371, 123 – 129, 1994.
McCave, I. N., Depositional features of organicrich black and green mudstones at DSDP Sites
386 and 387, western North Atlantic, Initial
Rep. Deep Sea Drill. Program, 43, 411 – 416,
1979.
McGillicuddy, D. J., Jr., and A. R. Robinson,
Eddy-induced nutrient supply and new production in the Sargasso Sea, Deep Sea Res., 44,1427 – 1450, 1997.
McGillicuddy, D. J., Jr., A. R. Robinson, D. A.
Siegel, H. W. Jannasch, R. Johnson, T. D.
Dickey, J. McNeil, A. F. Michaels, and A. H.
Knap, Influence of mesoscale eddies on new
production in the Sargasso Sea, Nature, 394,
263 – 266, 1998.
Menegatti, A. P., H. Weissert, R. S. Brown, R. V.
Tyson, P. Farrimond, A. Strasser, and M. Caron, High resolution d13C-stratigraphy throughthe early Aptian ‘‘Livello Selli’’ of the Alpine
Tethys, Paleoceanography, 13, 530 – 545,
1998.
Mutterlose, J., Temperature-controlled migrationof calcareous nannofloras in the northwest
European Aptian, in Nannofossils and Their
Applications, edited by J. A. Crux and S. E.
van Heck, pp. 122 – 142, Ellis Horwood, Chichester, UK, 1989.
Norris, R. D., Symbiosis as an evolutionary innovation in the radiation of Paleocene planktic
foraminifera, Paleobiology, 22, 461 – 480,
1996.
Norris, R. D., Recognition and macroevolutionary significance of photosymbiosis in molluscs, corals, and foraminifera, in Isotope
Paleobiology and Paleoecology, edited by W.
L. Manger and L. K. Meeks, Paleontol. Soc.
Pap., 4, 68 – 100, 1998.
Norris, R. D., and P. A. Wilson, Low-latitude
sea-surface temperatures for the mid-Cretaceous and the evolution of planktic foraminifera, Geology, 26, 823 – 826, 1998.
Norris, R. D., et al., Proceedings of the Ocean
Drilling Program Initial Report, vol. 171B,
Ocean Drill. Program, College Station, Tex.,
1998.
Opdyke, B. N., E. Erba, and R. L. Larson, Hot
LIPs, methane, and the carbon record of the
Apticore, Eos Trans. AGU, 80(46), F486 –F487, Fall Meet. Suppl., 1999.
Orth, C. J., M. Attrep, L. R. Quintana, W. P.

Elder, E. G. Kauffman, R. Diner, and T. Villamil, Elemental abundance anomalies in the late
Cenomanian extinction interval: A search forthe source(s), Earth Planet. Sci. Lett., 117,
189 – 204, 1993.
Oschlies, A., and V. Garcon, Eddy-induced enhancement of primary production in a modelof the North Atlantic Ocean, Nature, 394,
266 – 269, 1998.
Palmer, M. R., and G. G. Ernst, Generation ofhydrothermal megaplumes by cooling of pillow basalts at mid-ocean ridges, Nature, 393,
643 – 647, 1998.
Pearson, P. N., Stable isotopes and the study of
evolution in planktonic foraminifera, in Isotope Paleobiology and Paleoecology, edited
by W. L. Manger, and L. K. Meeks, Paleontol.
Soc. Pap., 4, 138 – 178, 1998.
Pearson, P. N., P. W. Ditchfield, J. Singano, K. G.
Harcourt-Brown, C. J. Nicholas, R. K. Olsson,
N. J. Shackleton, and M. A. Hall, Warm tropical sea surface temperatures in the Late Cretaceous and Eocene epochs, Nature, 413, 481 –
487, 2001.
Poulsen, C. J., E. J. Barron, C. C. Johnson, and P.
Fawcett, Links between major climatic factors
and regional oceanic circulation in the midCretaceous, in Evolution of the Cretaceous
Ocean-Climate System, edited by E. Barrera
and C. C. Johnson, Spec. Pap. Geol. Soc.
Am., 332, 73 – 89, 1999a.
Poulsen, C. J., E. J. Barron, W. H. Peterson, and
P. A. Wilson, A reinterpretation of mid-Cretaceous shallow marine temperatures throughmodel-data comparison, Paleoceanography,
14, 679 – 697, 1999b.
Poulsen, C. J., E. J. Barron, M. A. Arthur, andW. H. Peterson, Response of the mid-Cretaceous global ocean circulation to tectonic andCO2 forcings, Paleoceanography, 16, 576 –
592, 2001.
Pratt, L. M., and J. D. King, Low marine productivity and high eolian input recorded byrhythmic black shales in mid-Cretaceous pelagic deposits from central Italy, Paleoceanography, 1, 507 – 522, 1986.
Pratt, L. M., and C. N. Threlkeld, Stratigraphic
significance of 13C/12C ratios in mid-Cretaceous rocks of the Western Interior, U.S.A., in
The Mesozoic of Middle North America, edited
by D. F. Stott and D. J. Glass, Mem. Can. Soc.
Pet. Geol., 9, 305 – 312, 1984.
Pratt, L. M., M. A. Arthur, W. E. Dean, and P. A.
Scholle, Paleo-oceanographic cycles andevents during the Late Cretaceous in the Western Interior Seaway of North America, in Evolution of the Western Interior Basin, edited by
W. G. E. Caldwell and E. G. Kauffman, Spec.
Pap. Geol. Assoc. Can., 39, 333 – 354, 1993.
Premoli Silva, I., and W. V. Sliter, Cretaceousplanktonic foraminiferal biostratigraphy andevolutionary trends from the Bottaccione Section Gubbio, Italy, Palaeontogr. Ital., 81, 2 –90, 1995.
Premoli Silva, I., and W. V. Sliter, Cretaceous
paleoceanography: Evidence from planktonic
foraminiferal evolution, in Evolution of theCretaceous Ocean-Climate System, edited by
E. Barrera and C. C. Johnson, Spec. Pap. Geol.
Soc. Am., 332, 301 – 328, 1999.
Premoli Silva, I., E. Erba, and M. E. Tornaghi,
Paleoenvironmental signals and changes insurface fertility in mid-Cretaceous Corg-richpelagic facies of the fucoid marls (central
Italy), Geobios Mem. Spec., 11, 225 – 236,
1989.
Premoli Silva, I., E. Erba, G. Salvini, C. Locatelli, and D. Verga, Biotic changes in Cretac-

eous oceanic anoxic events of the Tethys, J.
Foraminiferal Res., 29, 352 – 370, 1999.
Pringle, M. S., and R. A. Duncan, Basement ages
from the southern and central Kerguelen Plateau: Initial products of the Kerguelen large
igneous province (abstract), Eos Trans. AGU,
81, Spring Meet. Suppl., abstract V31A-04,
2000.
Ravelo, A. C., and R. G. Fairbanks, Oxygenisotopic composition of multiple species ofplanktonic foraminifera: Records of the modern photic zone temperature gradient, Paleoceanography, 7, 815 – 831, 1992.
Ravelo, A. C., and R. G. Fairbanks, Carbon isotopic fractionation in multiple species ofplanktonic foraminifera from core-tops in thetropical Atlantic, J. Foraminiferal Res., 25,
53 – 74, 1995.
Retallack, G. J., and D. L. Dilcher, Cretaceous
angiosperm invasion of North America, Cretaceous Res., 7, 227 – 252, 1986.
Rich, J. E., G. L. Johnson, J. E. Jones, and J.
Campsie, A significant correlation between
fluctuations in seafloor spreading rates andevolutionary pulsations, Paleoceanography,
1, 85 – 95, 1986.
Richter, F. M., D. B. Rowley, and D. J. DePaolo,
Sr isotope evolution of seawater: The role oftectonics, Earth Planet. Sci. Lett., 109, 11 – 23,
1992.
Rivkin, R. B., et al., Vertical flux of biogenic
carbon in the ocean: Is there food web control?, Science, 272, 1163 – 1166, 1996.
Robaszynski, F., et al., Atlas de Foraminiféresplanctoniques du Crétacé moyen, parts 1-2,
Cah. Micropaleontol., 1 – 21979.
Roth, P. H., Mesozoic calcareous nannofossil
evolution: Relation to paleoceanographic
events, Paleoceanography, 2, 601 – 611, 1987.
Sageman, B. B., J. Rich, M. A. Arthur, W. E.
Dean, C. E. Savrda, and T. J. Bralower, Multiple Milankovitch cycles in the Bridge Creek
Limestone (Cenomanian-Turonian), WesternInterior Basin, in Stratigraphy and Paleoenvironments of the Cretaceous Western Interior
Seaway, USA, Concepts in Sedimentol. Paleontol., vol. 6, edited by W. E. Dean and
M. A. Arthur, pp. 153 – 171, Soc. Sed. Geol.,
Tulsa, Okla., 1998.
Schlanger, S. O., and H. C. Jenkyns, Cretaceous
oceanic anoxic events: Causes and consequences, Geol. Mijnbouw, 55, 179 – 184, 1976.
Schlanger, S. O., M. A. Arthur, H. C. Jenkyns,
and P. A. Scholle, The Cenomanian-Turonian
oceanic anoxic event, 1, Stratigraphy and distribution of organic carbon-rich beds and themarine C excursion, in Marine Petroleum
Source Rocks, edited by J. Brooks and A. J.
Fleet, Geol. Soc. Spec. Publ., 26, 371 – 399,
1987.
Schmidt, G. A., and L. A. Mysak, Can increased
poleward oceanic heat flux explain the warm
Cretaceous climate?, Paleoceanography, 11,579 – 593, 1996.
Scholle, P. A., and M. A. Arthur, Carbon isotope
fluctuations in Cretaceous pelagic limestones:
Potential stratigraphic and petroleum exploration tools, AAPGeol. Bull., 64, 67 – 87, 1980.
Shackleton, N. J., and J. P. Kennett, Paleotemperature history of the Cenozoic and the initiation of Antarctic glaciation: Oxygen andcarbon isotope analysis in DSDP Sites 277,
279, and 280, Initial Rep. Deep Sea Drill.
Proj., 29, 743 – 755, 1975.
Shipboard Scientific Party, Leg 183 summary:
Kerguelen Plateau-Broken Ridge—A large
igneous province, Proc. Ocean Drill. Program
Initial Rep., 183, 1 – 101, 2000.

LECKIE ET AL.: OCEANIC ANOXIC EVENTS AND PLANKTON EVOLUTION
Siegel, D. A., The Rossby rototiller, Nature, 409,
576 – 577, 2001.
Siegel, D. A., D. J. McGillicuddy Jr., and E. A.
Fields, Mesoscale eddies, satellite altimetry
and new production in the Sargasso Sea, J.
Geophys. Res., 104, 13,359 – 13,379, 1999.
Sigal, J., Chronostratigraphy and ecostratigraphyof Cretaceous formations recovered on DSDP
Leg 47B, Site 398, Initial Rep. Deep Sea Drill.
Proj., 47B, 287 – 327, 1979.
Signor, P. W., and G. J. Vermeij, The plankton
and the benthos: Origins and early history ofan evolving relationship, Paleobiology, 20,297 – 319, 1994.
Sikora, P. J., and R. K. Olsson, A paleoslopemodel of late Albian to early Turonian foraminifera of the western Atlantic margin andNorth Atlantic basin, Mar. Micropaleontol.,
18, 25 – 72, 1991.
Sinton, C. W., and R. A. Duncan, Potential links
between ocean plateau volcanism and global
ocean anoxia at the Cenomanian-Turonianboundary, Econ. Geol., 92, 836 – 842, 1997.
Sinton, C. W., R. A. Duncan, M. Storey, J. Lewis, and J. J. Estrada, An oceanic flood basalt
province within the Caribbean plate, Earth
Planet. Sci. Lett., 155, 221 – 235, 1998.
Sliter, W. V., Cretaceous foraminifers—Depth
habitats and their origin, Nature, 239, 514 –
515, 1972.
Sliter, W. V., Cretaceous foraminifers from thesouthwestern Atlantic Ocean, Leg 36, DeepSea Drilling Project, Initial Rep. Deep Sea
Drill. Proj., 36, 519 – 573, 1977.
Sliter, W. V., Mesozoic foraminifers and deep sea
benthic environments from Deep Sea Drilling
Project Sites 415 and 416, eastern North Atlantic, Initial Rep. Deep Sea Drill. Proj., 50,
353 – 370, 1980.
Sliter, W. V., Aptian anoxia in the Pacific basin,
Geology, 17, 909 – 912, 1989a.
Sliter, W. V., Biostratigraphic zonation for Cretaceous planktonic foraminifers examined inthin section, J. Foraminiferal Res., 19, 1 – 19,
1989b.
Sliter, W. V., Cretaceous planktic foraminiferal
biostratigraphy of the Calera Limestone, northern California, USA, J. Foraminifieral Res.,29, 318 – 339, 1999.
Snow, L. J., and R. A. Duncan, Hydrothermal
links between ocean plateau formation andglobal anoxia, Eos Trans. AGU, 82(47), Fall
Meet. Suppl., abstract OS41A-0437, 2001.
Southam, J. R., W. H. Peterson, and G. W. Brass,
Dynamics of anoxia, Palaeogeogr. Palaeoclimatol. Palaeoecol., 40, 183 – 198, 1982.
Stanley, S. M., and L. A. Hardie, Secular oscillations in the carbonate mineralogy of reefbuilding and sediment-producing organisms
driven by tectonically forced shifts in seawaterchemistry, Palaeogeogr. Palaeoclimatol. Palaeoecol., 144, 3 – 19, 1998.
Stoll, H. M., and D. P. Schrag, Evidence for
glacial control of rapid sea level changes inthe early Cretaceous, Science, 272, 1771 –
1774, 1996.
Stoll, H. M., and D. P. Schrag, High-resolution
stable isotope records from the Upper Cretaceous rocks of Italy and Spain: Glacial episodes
in a greenhouse planet?, Geol. Soc. Am. Bull.,
112, 308 – 319, 2000.
Sugarman, P. J., K. G. Miller, R. K. Olsson, J. V.
Browning, J. D. Wright, L. M. De Romero,
T. S. White, F. L. Muller, and J. Uptegrove,
The Cenomanian/Turonian carbon burial
event, Bass River, NJ, USA: Geochemical, pa-leoecological, and sea-level changes, J. Foraminiferal Res., 29, 438 – 452, 1999.
Summerhayes, C. P., Organic facies of middle
Cretaceous black shales in the deep North
Atlantic, AAPG Bull., 65, 2364 – 2380, 1981.
Summerhayes, C. P., Organic-rich Cretaceous sediments from the North Atlantic, in Marine
Petroleum Source Rocks, edited by J. Brooks
and A. J. Fleet, Geol. Soc. Spec. Publ., 26,
301 – 316, 1987.
Tappan, H., Phytoplankton: Below the salt at the
global table, J. Paleontol., 60, 545 – 554, 1986.
Tappan, H., and A. R. Loeblich Jr., Evolution of
the oceanic plankton, Earth Sci. Rev., 9, 207 –
240, 1973.
Tarduno, J. A., W. V. Sliter, L. Kroenke, M.
Leckie, H. Mayer, J. J. Mahoney, R. Musgrave, M. Storey, and E. L. Winterer, Rapid
formation of Ontong Java Plateau by Aptian
Mantle Plume Volcanism, Science, 254,
399 – 403, 1991.
Tejada, M. L. G., J. J. Mahoney, R. A. Duncan,
and M. P. Hawkins, Age and geochemistry ofbasement and alkalic rocks of Malaita and Santa Isabel, Solomon Islands, southern margin of
Ontong Java Plateau, J. Petrol., 37, 361 – 394,
1996.
Theyer, C. W., Sediment-mediated biologicaldisturbance and the evolution of marine
benthos, in Biotic Interactions in Recent andFossil Benthic Communities, edited by M. J. S.
Tevesz and P. L. McCall, pp. 479 – 625, Plenum, New York, 1983.
Thomas, E., and N. J. Shackleton, The Paleocene-Eocene benthic foraminiferal extinction
and stable isotope anomalies, in Correlationof the Early Paleogene in Northwest Europe,edited by R. W. O. Knox and R. E. Dunay,
Geol. Soc. Spec. Publ., 101, 401 – 441, 1996.
Thurow, J., H.-J. Brumsack, J. Rullkötter, R.
Littke, and P. Meyers, The Cenomanian/Turonian boundary event in the Indian Ocean—A
key to understand the global picture, in Synthesis of Results From Scientific Drilling in the
Indian Ocean, Geophys. Monogr. Ser., vol. 70,
pp. 253 – 273, AGU, Washington, D.C., 1992.
Tjalsma, R. C., and G. P. Lohmann, PaleoceneEocene bathyal and abyssal benthic foraminifera from the Atlantic Ocean, Micropaleontol.
Spec. Publ., 4, 1 – 90, 1983.
Toggweiler, J. R., An ultimate limiting nutrient,
Nature, 400, 511 – 512, 1999.
Tornaghi, M. E., I. Premoli Silva, and M. Ripepe,
Lithostratigraphy and planktonic foraminiferal
biostratigraphy of the Aptian-Albian ‘‘Scisti a
Fucoidi’’ in the Piobbico core, Marche, Italy:
Background for cyclostratigraphy, Riv. Ital.
Paleontol. Strat., 95, 223 – 264, 1989.
Tortell, P. D., M. T. Maldonado, and N. M. Price,
The role of heterotrophic bacteria in iron-limited ocean ecosystems, Nature, 383, 330 – 332,
1996.
Tucholke, B. E., and P. R. Vogt, Western North
Atlantic: Sedimentary evolution and aspects oftectonic history, Initial Rep. Deep Sea Drill.
Proj., 43, 791 – 825, 1979.
Tyrell, T., The relative influences of nitrogen andphosphorous on oceanic primary productivity,
Nature, 400, 525 – 531, 1999.
Uz, B. M., J. A. Yoder, and V. Osychny, Pumping nutrients to ocean surface waters by theaction of propagating planetary waves, Nature,409, 597 – 600, 2001.
Vermeij, G. J., The Mesozoic marine revolution:
Evidence from snails, predators, and grazers,
Paleobiology, 3, 245 – 258, 1977.

13 - 29

Vermeij, G. J., Economics, volcanoes, and Phanerozoic revolutions, Paleobiology, 21, 125 –
152, 1995.
Vogt, P. R., Volcanogenic upwelling of anoxic
nutrient-rich water: A possible factor in carbonate-bank/reef demise and benthic faunal extinctions, Geol. Soc. Am. Bull., 101, 1225 –1245, 1989.
Wei, K.-Y., and J. P. Kennett, Taxonomic evolution of Neogene planktonic foraminifera andpaleoceanographic relations, Paleoceanography, 1, 67 – 84, 1986.
Weissert, H., and A. Lini, Ice age interludes during the time of Cretaceous greenhouse climate?, in Controversies in Modern Geology,
edited by D. W. Mueller et al., pp. 173 – 191,
Academic, San Diego, Calif., 1991.
Weissert, H., A. Lini, K. B. Föllmi, and O. Kuhn,
Correlation of Early Cretaceous carbon isotope
stratigraphy and platform drowning events: A
possible link?, Palaeogeogr. Palaeoclimatol.
Palaeoecol., 137, 189 – 203, 1998.
Whitechurch, H., R. Montigny, J. Sevigny, M.
Storey, and V. Salters, K-Ar and 40Ar-39Ar
ages of central Kerguelen Plateau basalts,
Proc. Ocean Drill. Program Sci. Results,
120, 71 – 77, 1992.
Wilde, P., and W. B. N. Berry, Progressive ventilation of the oceans: Potential for return toanoxic conditions in the post-Paleozoic, in
Nature and Origin of Cretaceous Carbon-Rich
Facies, edited by S. Schlanger and M. B. Cita,
pp. 209 – 224, Academic, San Diego, Calif.,
1982.
Williams, R. G., and M. J. Follows, Eddies makeocean deserts bloom, Nature, 394, 228 – 229,
1998.
Wilson, P. A., and R. D. Norris, Warm tropicalocean surface and global anoxia during the
mid-Cretaceous period, Nature, 412, 425 –
429, 2001.
Wonders, A. A. H., Middle and Late Cretaceousplanktonic foraminifera of the western Mediterranean area, Utrecht Micropaleontol. Bull.,
24, 1 – 157, 1980.
Woo, K. S., T. F. Anderson, L. B. Railsback, and
P. A. Sandberg, Oxygen isotope evidence forhigh salinity surface seawater in the mid-Cretaceous Gulf of Mexico: Implications for warm
saline deepwater formation, Paleoceanography, 7, 673 – 685, 1992.
Zachos, J., L. D. Stott, and K. C. Lohmann,
Evolution of early Cenozoic marine temperatures, Paleoceanography, 9, 353 – 387, 1994.
Zeebe, R. E., Seawater pH and isotopic paleotemperatures of Cretaceous oceans, Palaeogeogr. Palaeoclimatol. Palaeoecol., 170, 49 –57, 2001.
Zimmerman, H. B., A. Boersma, and F. W.
McCoy, Carbonaceous sediments and paleoenvironment of the Cretaceous South Atlantic
Ocean, in Marine Petroleum Source Rocks,
edited by J. Brooks and A. J. Fleet, Geol.
Soc. Spec. Publ., 24, 271 – 286, 1987.

T. J. Bralower, Department of Geological
Sciences, Campus Box 3315, University ofNorth Carolina, Chapel Hill, NC 27599-3315,
USA.
R. Cashman and R. M. Leckie, Department ofGeosciences, University of Massachusetts, 611
North Pleasant Street, Amherst, MA 01003,
USA. (mleckie@geo.umass.edu)