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Neogene overflow of Northern Component Water at theGreenland-Scotland Ridge

H. R. PooreDepartment of Earth Sciences, University of Cambridge, Downing Street, Cambridge CB2 3EQ, UK([email protected])

R. SamworthStatistical Laboratory, Centre for Mathematical Studies, University of Cambridge, Wilberforce Road, Cambridge, CB30WB, UK

N. J. WhiteDepartment of Earth Sciences, University of Cambridge, Downing Street, Cambridge CB2 3EQ, UK

S. M. JonesDepartment of Geology, Trinity College, University of Dublin, Dublin 2, UK

I. N. McCaveDepartment of Earth Sciences, University of Cambridge, Downing Street, Cambridge CB2 3EQ, UK

[1] In the North Atlantic Ocean, flow of North Atlantic Deep Water (NADW), and of its ancientcounterpart Northern Component Water (NCW), across the Greenland-Scotland Ridge (GSR) is thought tohave played an important role in ocean circulation. Over the last 60 Ma, the Iceland Plume has dynamicallysupported an area which encompasses the GSR. Consequently, bathymetry of the GSR has varied withtime due to a combination of lithospheric plate cooling and fluctuations in the temperature and buoyancywithin the underlying convecting mantle. Here, we reassess the importance of plate cooling and convectivecontrol on this northern gateway for NCW flow during the Neogene period, following Wright and Miller(1996). To tackle the problem, benthic foraminiferal isotope data sets have been assembled to examined13C gradients between the three major deep water masses (i.e., Northern Component Water, SouthernOcean Water, and Pacific Ocean Water). Composite records are reported on an astronomical timescale, anda nonparametric curve-fitting technique is used to produce regional estimates of d13C for each water mass.Confidence bands were calculated, and error propagation techniques used to estimate %NCW and itsuncertainty. Despite obvious reservations about using long-term variations of d13C from disparate analysesand settings, and despite considerable uncertainties in our understanding of ancient oceanic transportpathways, the variation of NCW through time is consistent with independent estimates of the temporalvariation of dynamical support associated with the Iceland Plume. Prior to 12 Ma, d13C patterns overlapand %NCW cannot be isolated. Significant long-period variations are evident, which are consistent withpreviously published work. From 12 Ma, when lithospheric cooling probably caused the GSR to submergecompletely, long-period d13C patterns diverge significantly and allow reasonable %NCW estimates to bemade. Our most robust result is a dramatic increase in NCW overflow between 6 and 2 Ma whendynamical support generated by the Iceland Plume was weakest. Between 6 and 12 Ma a series ofvariations in NCW overflow have been resolved.

Components: 13,280 words, 12 figures, 3 tables.

Keywords: Northern Component Water; Greenland-Scotland Ridge; Iceland; plume; Neogene; d13C.

G3G3GeochemistryGeophysics

Geosystems

Published by AGU and the Geochemical Society

AN ELECTRONIC JOURNAL OF THE EARTH SCIENCES

GeochemistryGeophysics

Geosystems

Article

Volume 7, Number 6

14 June 2006

Q06010, doi:10.1029/2005GC001085

ISSN: 1525-2027

Copyright 2006 by the American Geophysical Union 1 of 24

Index Terms: 3037 Marine Geology and Geophysics: Oceanic hotspots and intraplate volcanism; 3275 Mathematical

Geophysics: Uncertainty quantification (1873); 4850 Oceanography: Biological and Chemical: Marine organic chemistry

(0470, 1050).

Received 22 July 2005; Revised 6 March 2006; Accepted 5 April 2006; Published 14 June 2006.

Poore, H. R., R. Samworth, N. J. White, S. M. Jones, and I. N. McCave (2006), Neogene overflow of Northern Component

Water at the Greenland-Scotland Ridge, Geochem. Geophys. Geosyst., 7, Q06010, doi:10.1029/2005GC001085.

1. Introduction

[2] Wright and Miller [1996] argue that there is astrong negative correlation between the percentageof Northern Component Water (%NCW), whichaffects d13C values in the Southern Ocean, andvertical motions of the Greenland-Scotland Ridge(GSR) during Neogene times (Figures 1–3). Thesevertical motions occur with at least two timescales.First, long-term subsidence of the GSR (timescaleof 10s of millions of years) is controlled by coolingof the lithospheric plate as it drifts away from themid-oceanic ridge. Secondly, short-term uplift andsubsidence excursions (timescale of 1–10 Ma) arecontrolled by rapid buoyancy fluctuations withinthe Iceland Plume. This plume is a major convec-tive upwelling within the mantle which has had animportant effect upon the structural development ofthe North Atlantic Ocean over the last 60 Ma.

[3] Such changes in elevation of the GSR couldaffect flux of NCW into the Southern and PacificOceans. When the GSR is deeper (i.e., the IcelandPlume is cooler) a larger volume of deep water canflow over the ridge, thus increasing %NCW pres-ent in the Southern Ocean (Figure 3). When theIceland Plume is hotter and the GSR is shallower,production of NCW could be significantly reduced.This lock-gate mechanism is important because itmight help to improve our understanding of tem-poral changes within the Iceland Plume and thuscontribute to knowledge of the dynamics of mantleconvection. It may also help to gauge the strengthof thermohaline circulation and hence the influenceof variations in NCW production on Neogeneclimate.

[4] Present-day NCW (i.e., North Atlantic DeepWater, NADW) has three principal sources(Figure 1): (1) overspill from the Norwegian-Greenland Seas through the Faroe-Shetland Chan-nel (1.7 Sv); (2) overspill through the DenmarkStraits (2.9 Sv); and (3) convective overturningwithin the Labrador Sea [Dickson and Brown,1994]. It remains unclear when the Labrador Sea

became an important source of deep water[Mikolajewicz and Crowley, 1997; Burton et al.,1997; Lear et al., 2003]. Roughly 40% of the totalNADW flux is sourced from the Labrador Sea[Schmitz and McCartney, 1993]. We assume thatif the cold deep core of NADW from the Faroe-Shetland Channel and from the Denmark Straits isreduced significantly, then a large reduction in%NCW in the Southern Ocean results becausethe cold deep core of NADW consists primarilyof GSR overflow water and not of warmer, lesssalty Labrador Sea Water.

[5] At present, flow of cold, deep NCW from theNorwegian-Greenland Sea into the North AtlanticOcean is impeded by the GSR, which extends eastand west from Iceland (Figure 1b). The present-daydepth of the deeper channels through this ridge isno greater than �850 m and allows a total of 5.6 Sv(106 m3) of deep water to flow over the barrier[Dickson and Brown, 1994], although its volumeincreases by entrainment downstream of channelexits. Hence Norwegian Sea Overflow Water(NSOW) forms much of the Lower North AtlanticDeep Water (LNADW). At present, the core ofNSOW flows at depths of 500–600 m, descendingas it travels southward through the Faroe-BankChannel [Kuijpers et al., 2002]. Cold northernwaters are restricted to �150 m above the seabed in this deepest gateway (Figure 1b). ThereforeGSR overflow waters should be sensitive to plume-related dynamic uplift.

[6] Over the last 10 years, much new benthicforaminiferal d13C (and d18O) data have becomeavailable. These new data have an average sampleinterval of up to �5 ka and facilitate improvedcore-to-core correlation (Figures 4–6) [Billups etal., 2004]. North Atlantic records include data fromODP Site 925 [Billups et al., 1997, 1998; H. Pfuhl,written communication, 2004]. Pacific Oceanrecords include analyses from ODP Site 846 (0–7.4 Ma [Mix and Shackleton, 1995; Shackleton etal., 1995b]). Most of the Southern Ocean data isfrom ODP Legs 177 and 185. The geological

GeochemistryGeophysicsGeosystems G3G3

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timescale has also undergone revision thanks to theavailability of astronomically tuned records for theNeogene period [Lourens et al., 2004]. With agreatly improved database and timescale, we canmore accurately construct end-member records ofcarbon isotopes. Finally, we present a new strategyfor fitting d13C records with smooth functions andconfidence bands. In this way, we can estimate thevariation of NCW overflow and its associated

uncertainty through time and revisit the Wrightand Miller [1996] hypothesis.

2. Benthic Carbon Isotope Data Sets

[7] A range of proxies is used to understandpaleocirculation patterns, including d13C and tracemetals (e.g., Neodymium, Cadmium). FollowingWright and Miller [1996], we exploit d13C because

Figure 1. (a) Bathymetric map of North Atlantic Ocean. Principal features include Greenland-Scotland Ridge(GSR), Iceland-Faroe Ridge (IFR), Faroe-Shetland Channel (FSC), Rockall Trough (RT), Denmark Straits (DS),Norwegian-Greenland Sea (NGS), and Irminger Basin (IrB). V-shaped ridges (VSRs) are prominent to the southwestof Iceland (see Figure 2). Arrows show deep-water currents which overflow GSR. Fluxes are in Sverdrups (1 Sv =1 � 106 m3) beside respective arrow [Dickson and Brown, 1994]. A–A0 shows location of section. (b) Schematicsection across Greenland-Scotland Ridge redrawn from Miller and Tucholke [1983] and Wright and Miller [1996].Depths given for major spill-points of NCW over GSR.


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samples can be easily assigned ages, records are ofsufficiently high resolution, and there is now awealth of published and unpublished analyses.d13C is a useful tracer of water masses becauseit reflects the nutrient content of deep water[Ruddiman, 2001], although care must be takenin choosing species that accurately record theisotopic composition of seawater. North Atlantic

Deep Water is 13C rich because surface waters fromwhich it forms have previously been depleted inboth nutrients and 12C (manifested by a decrease of1%) within the subtropical Atlantic Ocean. Incontrast, deep water formed in the Weddell Seacontains more nutrients and less 13C becausesurface waters sink without being reventilated.d13C is not a conservative tracer. However, givena residence time of deep water in the deep Atlanticof �300 yrs [Broecker and Peng, 1982], it isreasonable to model benthic d13C as a conservativetracer. Benthic d13C can then be used to quantifythe mixing of two water masses at a particular site[Kroopnick, 1985; Oppo and Fairbanks, 1987].One minor complication in understanding benthicd13C gradients is the existence of the ‘‘Mackenseneffect’’ [Mackensen et al., 1993], which can lead toextremely low d13C values. However, this effect isprobably smaller than the ventilation effect and asfar as is possible we use epibenthic foraminiferasuch as Cibicidoides wuellerstorfi, which shouldcircumvent this problem.

2.1. Benthic D13C End-Members

[8] Benthic stable isotope data have been compiledfrom sediment cores located in the North AtlanticOcean, in the Atlantic sector of the Southern Oceanand in the Pacific Ocean (Figures 4 and 5). Eachregion acts as an end-member in the mixingequation of Oppo and Fairbanks [1987] which isused to find an estimate of the percentage overflow

Figure 2. Residual free-air gravity map of NorthAtlantic Ocean after Jones et al. [2002a]. Gray lozengeshows approximate location of Greenland-ScotlandRidge; numbered boxes show identified V-shaped ridgesreferred to in text and in Figure 9.

Figure 3. (a) Percentage overflow of Northern Component Water (%NCW). (b) Residual depth data. Both takenfrom Wright and Miller [1996]. Note negative correlation between %NCW in the Southern Ocean through time andresidual depth of oceanic crust at Iceland.



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of NCW through time. The estimate of %NCW,dNCW, is given by

dNCW ¼ d13CSO � d13CPO

d13CNA � d13CPO

� 100; ð1Þ

where d13CSO, d13CPO and d13CNA are estimates of

d13C for the Southern, Pacific, and North AtlanticOceans. Note that in the text, we refer to %NCWrather than its estimate.

[9] It is important to choose the geographicallocation of each data set carefully so that the truevariation of each end-member is reflected in thecomposite record. North Atlantic records are takenfrom sites which are sufficiently far north and fromdepths at which benthic foraminifera are likely tobe fully bathed by the North Atlantic end-member.Southern Ocean sites should ideally be locatedbeneath the mixing zone of Upper CircumpolarDeep Water (UCDW) and NCW, a boundary thatwe assume remains fixed through time. This mix-ing zone is likely to have changed both its geo-graphical position and its depth, but we lack data toconstrain any movement during Neogene times.Pacific Ocean sites should reflect Pacific Oceandeep water and therefore should not includerecords from restricted sub-basins such as ODPSite 677 [Mix and Shackleton, 1995; Oppo et al.,1995]. The Pacific Ocean can exhibit east-westd13C gradients which result from nonconservative

d13C effects (i.e., water-mass aging caused byaccumulation of nutrients [Kroopnick, 1985]). Pa-cific d13C gradients are problematic but difficult tocircumvent, given the geographical distribution ofreliable Neogene high-resolution data [Whitmanand Berger, 1993].

[10] Sources of data used for each oceanic end-member are given in Figure 5 and in Tables 1–3.d13C data are quoted to the PDB standard andspecies disequilibria are corrected using the speciesoffsets of Graham et al. [1981], Shackleton andHall [1984], Muller et al. [1991], and Katz et al.[2003]. Where multiple samples are used or whereindividual data points are obviously outliers, d13Cvalues obtained from more unreliable species havebeen removed [Katz et al., 2003]. We have alsoremoved data points for which both carbon andoxygen isotopes were greater than two estimatedstandard deviations from the regression curve. Wesubsequently refitted the data using our local linearkernel regression procedure. The distribution ofstandard deviations then better approximated anormal distribution, which is an assumption madeduring the fitting procedure.

[11] Records used to construct Atlantic and Pacificend-members are mostly similar to those used byWright and Miller [1996]. Our North Atlanticcomposite record now includes ODP Site 925 forpart of the late Miocene and Early Pliocene

Figure 4. Map showing thermohaline conveyor belt [Rahmstorf, 2002]. Solid black lines, transport pathways ofbottom water; dark gray lines, transport pathways of cold, deep water; light gray lines, transport pathways of warm,surface water; ellipses, locations of major downwelling zones; solid black circles, sites of DSDP and ODP cores usedin this study.



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(6.240–3.281 Ma). Billups et al. [1998] haveshown that Site 925, at 3042 m mostly reflectsNADW which is dominant to depths of 3000 m onCeara Rise during the Early Pliocene. Despite itssoutherly location we therefore think that Site 925adequately records the variation in d13C of NCW.In constructing the Pacific end-member, we havealso replaced ODP Site 677 with ODP Site 846 and

have supplemented DSDP Site 289 with data fromDSDP Sites 296, 574A and ODP Sites 806, 846.Our most significant revision concerns the South-ern Ocean end-member. Recent ODP cruises haverecovered sediment cores which span the ageinterval of interest and which lie beneath themodern mixing zone of CDWand NADW [Billups,2002]. New material from Leg 177 Sites 1088

Figure 5. Assemblage of d13C isotopic records determined from benthic foraminifera. (a) North Atlantic Ocean.(b) Southern Ocean. (c) Pacific Ocean. Each record is color-coded according to site (see Tables 1–3 for furtherdetails). Box shows location of Figure 14.



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[Billups, 2002], and 1090 [Billups et al., 2002,2004] provide an opportunity to reexamine theSouthern Ocean composite d13C record in detail.

2.2. Age Models

[12] Accurate dating of compiled records is a majorconcern in order to ensure that d13C gradients canbe reliably interpreted. Astronomical dating ofmagnetobiostratigraphic datums significantlyimproves the correlation of different records. d13Crecords from each site were placed on the astro-nomically calibrated timescale of Lourens et al.[2004] by linearly interpolating between magneto-biostratigraphic datums (Figure 5). Age models forthe different benthic d18O records were then ad-

justed by comparison with an astronomically tunedd18O record in order to refine correlation of thethree oceanic end-members (Figure 6). A compos-ite isotope sequence from V19–30, ODP 677 andODP 846 was used for the 0–6 Ma interval (http://delphi.esc.cam.ac.uk [Shackleton and Pisias, 1985;Shackleton et al., 1995a, 1990, 1995b]). Data fromODP Site 982 [Hodell et al., 2001] were usedbetween 6 and 9 Ma. A sequence from ODP Site926 was used for the 14–16 Ma interval (N.J.Shackleton, written communication 2004) and datafrom ODP Site 1090 were used to calibrate the restof the Neogene reference d18O record [Billups etal., 2004]. The 9–14 Ma interval does not havehigh resolution benthic d18O records from siteswhich have been astronomically tuned. Instead, a

Table 1. Sources of Data Used in North Atlantic d13C End-Member

Site Isotope Source Age Model Source Interval, Ma No. Points

552A Shackleton and Hall [1984] Shackleton and Hall [1984] 0–3.446 604Keigwin et al. [1986] 5.617–8.845 418Curry and Miller [1989] Curry and Miller [1989] 2.203–2.519 36

2.810–2.872Keigwin [1984] Shackleton and Hall [1984] 7–8.271 4

Keigwin et al. [1986]563 Miller and Fairbanks [1985] Miller and Fairbanks [1985] 8.995–27.965 59

Wright et al. [1991] Miller and Fairbanks [1985] 9.223–11.121 39Wright et al. [1992] Miller and Fairbanks [1985] 11.154–23.377 97

607 Ruddiman et al. [1989] Ruddiman et al. [1989] 0.4–0.451 119Clement and Robinson [1986] 0.522–0.567

0.750–0.8101.251–1.648

Raymo et al. [1992] Clement and Robinson [1986] 3.152–3.200 16608 Wright et al. [1991] Clement and Robinson [1986] 7.994–9.113 11

Miller et al. [1991] 11.020–11.210Wright et al. [1992] Clement and Robinson [1986] 11.301–11.422 48

Miller et al. [1991] 16.466–21.915Miller et al. [1987] Clement and Robinson [1986] 6.618–9.000 32

12.414–21.872925 [Billups et al., 1997] Shackleton and Crowhurst [1997] 3.281–6.240 339

Pfuhl (unpub.) Tiedemann and Franz [1997] 3.281–6.240 93Pfuhl and Shackleton [2004] 5.0–6.240 259

Table 2. Sources of Data Used in Southern Ocean d13C End-Member


360 Wright et al. [1992] Jenkins [1978] 12.413–13.922 9704 Hodell and Ciesielski [1991] Hailwood and Clement [1991] 0.056–5.377 642

Hodell and Ciesielski [1991]Wright et al. [1991] Hailwood and Clement [1991] 9.155–9.632 17

10.920–11.890747 Wright and Miller [1992] Heider et al. [1992] 11.005–16.065 57

Berggren [1992] 24.12–26.37Billups (unpub. data) as above 0.667–25.374 27

1088 Billups [2002] Shipboard Scientific Party [1999] 0–11.004 1581090 Billups et al. [2004] Billups et al. [2004] 15.94–24.19 1532



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combination of Leg 154 magnetic susceptibilityand reflectance data together with d18O recordsfrom sites 704 and 747, which have reliable mag-netostratigraphy and continuous sedimentation,were used as reference sequences to which otherdata sets could be correlated. The age model for 9–14 Ma does have large uncertainties compared withdata outside this range but the d13C records aresparse, so correlation to future astronomicallytuned benthic oxygen data is unlikely to alter theage models significantly. When depth-age con-straints were added to the overall age model,resulting sedimentation rates were checked to en-sure that sedimentation rates varied smoothly ex-cept at known hiatuses and unconformities.

2.2.1. Spectral Analysis

[13] Spectral analysis of the corresponding benthicoxygen isotope data provides an useful check onthe quality of age models applied to the carbonisotope data. The oxygen isotopic composition ofsea water primarily reflects global ice volume[Imbrie et al., 1984], which responds to changesin the seasonal input of insolation at differentlatitudes. These changes are controlled by threeorbital parameters of the Earth: eccentricity (100and 400 ka periods), obliquity (41 ka) and preces-sion (19 and 23 ka).

[14] Both Southern and Pacific Oceans show thatsignificant orbital frequencies are present, mostlyas a result of the inclusion of astronomically tuneddata sets. All oxygen data are coherent with thecalculated insolation curve for higher orbital perio-dicities in both these oceans. Precession frequen-cies are only weakly coherent (below 80%significance level) suggesting that data are poorlytuned at these periodicities. This weakness isunsurprising given the caveats mentioned above,the temporal range of the data, and the coarsersampling intervals for the Miocene epoch com-

pared with Plio-Pleistocene times. North Atlanticcomposite records have the weakest periodicitieswith peaks in power which are barely significantcompared with background noise.

3. Statistical Analysis

[15] There are significant errors in d13C measure-ments and in their ages, both of which propagatethrough the differences in equation (1) to affect ourestimate of %NCW. The source of error in d13C isboth measurement error, whose standard deviationranges between 0.05 and 0.1%, and in vital effectspecies correction factors, which [Graham et al.,1981] quote as ranging between 0.12 and 0.46%.In order to calculate the variation of %NCW withtime, each composite d13C record is first fitted witha smooth function. There are numerous ways inwhich this fitting can be carried out. Here, weoutline an approach which has two importantfeatures. First, we allow the bandwidth of thekernel estimator to vary through time. This flexi-bility ensures that large variations in sample distri-bution are accommodated. Secondly, confidencebands for best-fitting functions are constructed.We give a brief discussion of the techniques usedand refer the interested reader to Samworth andPoore [2005] for further details.

3.1. Local Linear Kernel Regression

[16] Our first task is to fit a smooth curve througheach composite record. Gaussian filters are oftenused for this purpose [e.g., Mix and Shackleton,1995; Wright and Miller, 1996; Zachos et al.,2001]. In the statistical literature, this locallyweighted mean would be referred to as a localconstant kernel estimator with a Gaussian kernel orweight function. The more sophisticated locallinear kernel estimator, however, has the advantageof significantly reduced bias, particularly at the

Table 3. Sources of Data Used in Pacific Ocean d13C End-Member


289 Savin et al. [1981] Srinivasan and Kennett [1981] 4.797–21.521 220Woodruff and Savin [1989] Shipboard Scientific Party [1975] 5.538–22.060 12Woodruff and Savin [1991] 13.203–21.896 42

296 Savin et al. [1981] Ellis [1975] 6.099–11.454 1617.864–20.834 30

574A Pisias et al. [1985] Gartner and Chow [1985] 11.604–16.012 495806 Corfield and Cartlidge [1993] Takayama [1993] 11.951–18.881 126846 Mix and Shackleton [1995] Shackleton et al. [1995a] 0–1.849 697846 Shackleton et al. [1995b] Shackleton et al. [1995a] 1.763–7.644 23141170 Pfuhl and McCave [2004] Pfuhl and McCave [2004] 20.119–28.112 269



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Figure 6. Assemblage of d18O records determined from benthic foraminifera. (a) Records used as reference model.Records from (b) North Atlantic, (c) Southern, and (d) Pacific Oceans which were used to refine age models.



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boundaries of the data, or in regions of high sampleresolution [Fan and Gijbels, 1996]. The locallinear estimator for (xi, yi) pairs is given by

mh xð Þ ¼ 1

n

Xni¼1

s2 � s1 xi � xð Þf gKh xi � xð Þs2s0 � s21

yi ð2Þ

where

sr ¼1

n

Xni¼1

xi � xð ÞrKh xi � xð Þ; r ¼ 0; 1; 2;

and the scaled kernel, Kh(�) = h�1K(�/h), satisfiesRKh = 1.

[17] For both local constant and local linear esti-mators, the crucial choice is that of the bandwidth,h, which controls the amount of smoothing: toosmall a value of h results in a spiky, under-smoothed curve; equally if h is too large then thedata are oversmoothed leading to a loss of thesignal.

[18] Oceanographic records are generally fittedusing a Gaussian filter with a constant bandwidth.The main difficulty with this approach is that datadistribution is inevitably highly nonuniform as aconsequence of records being pooled from differ-ent cores, sampled at different resolutions, andbecause of gaps and unconformities. A bandwidthwhich is appropriate in a region of high resolutiondata may be many times smaller than that in aregion of sparse data. We have sidestepped thisissue by dividing the age range into blocks ofapproximately equal sample resolution; we thenchose a separate bandwidth for each block. Sincewe are less interested in the rate of change of d13C,continuity of first and second derivatives is not ofprimary concern. Pacific, Southern and North At-lantic composite records were divided into 21, 16and 12 blocks, respectively, and appropriate band-widths were chosen for each block. Our regressionmethod slides the chosen kernel through the agerange of the composite record. At boundaries ofblocks, part of the kernel cannot be used to weight(xi, yi) pairs, which results in a large bias. To reducethis effect, we take the half-width of the kernel andextend the age range of each block by this width.

3.2. SIMEX Algorithm

[19] In addition to the error in the y-directionassociated with measuring d13C level at a givenage, there is also an error in the x-direction asso-ciated with assigning an age to a particular sample.This error arises from several sources, including

the assumption of a constant sedimentation ratebetween age-depth control points and the fact thatthe ages and composite depths of these controlpoints are not precisely known. Temporal errorshave often been ignored in previous studies. How-ever, errors-in-variables techniques have been de-veloped in statistical analysis for handling thisproblem.

[20] We have applied the Simulation Extrapolation(SIMEX) algorithm of Cook and Stefanski [1994]and Carroll et al. [1995] to all data. The results forthe first block of the Pacific Ocean record areshown in Figure 7. The SIMEX algorithm analysesthe effect of adding additional noise of knownstandard deviation to the age of each point and itthen extrapolates backward to estimate the effect ofhaving no noise at all. Addition of this extra noisetends to blur peaks and troughs in the estimatedcurve. The final SIMEX estimate helps to clarifythese features.

3.2.1. Confidence Bands

[21] We have calculated 95% confidence bands foreach block in order to estimate uncertainties ofeach composite d13C record. Each confidence bandis interpreted as a region which contains the true(i.e., unobserved) d13C variation over the age rangeof the block in question with a probability of 0.95.Many techniques for the construction of confidencebands in nonparametric regression problems havebeen published in the statistical literature. Welooked in detail at those proposed by Xia [1998]and at two constructions proposed by Claeskensand van Keilegom [2003]. We prefer the Xia [1998]confidence bands which are generally wider thanthose of the first construction of Claeskens and vanKeilegom [2003]. Their second construction wasfound to be very sensitive to small changes inbandwidth for our data. The Xia [1998] bandsbehave in a more robust and conservative waywhich is suitable for our purposes (Figure 8). Thesebands also allow bias of the regression estimate atpeaks and troughs to be estimated so that bands arenot simply centered around the regression estimate.Unfortunately, there is currently no theoreticalsupport for calculating confidence bands aroundSIMEX estimates. Therefore we only calculate Xiabands around the bias-corrected local linear regres-sion estimate.

[22] Confidence bands for blocks with smallamounts of data (usually blocks with a small agerange) must be treated with caution because ofundercoverage problems. In other words, the prob-



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ability that a confidence band contains the trueregression curve may be less than 95% if there is alimited amount of data (i.e., �50 data points perblock). Intervals that may be particularly suscepti-ble to this problem lie between 6 and 7 Ma, and12.5 and 13.5 Ma in the Pacific Ocean.

4. End-Member Regression Estimates

[23] Figure 8 shows regression estimates for thethree end-members over the last 23 Ma. The widthof each 95% confidence band suggests that we canconsider these composite records in two parts.Before �12 Ma, all three composite records areeffectively identical with the same pattern occur-ring in each ocean. After �12 Ma, the threecomposite records diverge by up to about 1.5%.

[24] In general terms, our results confirm well-known behavior. Circumpolar Deep Water(CPDW) represents a mixture of NADW andrecirculated deep water from the Indian and Pacific

Oceans. Consequently, CPDW d13C values liebetween those of NADW (characterized by highd13C and low nutrients) and Indo-Pacific DeepWater (low d13C and high nutrients). Without anyNADW influence, the d13C of CPDW shoulddecrease to Indo-Pacific values, regardless of howthe oceanic inventory of 12C or 13C might havechanged [Charles and Fairbanks, 1992]. Thus theobservation that Atlantic, Southern and PacificOcean d13C end-member records coincide prior to�12 Ma suggests that significant NCW productionbegan at about 12 Ma. Detailed age calibrationand error analysis strengthens earlier suggestionsbased on carbon isotopes that NCW production wasweak during middle Miocene times (16–11.6 Ma)but became more important after about 12 Ma[Woodruff and Savin, 1989, 1991; Wright et al.,1992; Wright and Miller, 1993]. The magnitude ofa d13C gradient, or the lack of any such gradient,may also be due to lower whole ocean nutrients[e.g., Delaney and Boyle, 1987] and may be thecause of the similarities between the North Atlantic,Southern and Pacific Ocean composite recordsduring early and middle Miocene times.

[25] Changes in the global carbon reservoir arethought to cause long-period variations in d13C[Miller and Fairbanks, 1985; Vincent and Berger,1985; Shackleton, 1987]. These reservoir-relatedd13C variations are most noticeable in the period23–12 Ma. For example, the Monterey Excursionbetween �17.5 Ma and 13.5 Ma is a well-knownfeature of Miocene benthic d13C records which isthought to be caused by a period of increasedorganic carbon burial [Vincent and Berger, 1985].d13C also decreases between �22 and 20 Ma,which may correlate with shallowing of the CCDidentified by van Andel [1975] and van Andel et al.[1975]. Lyle [2003] attributes the shallowing of theCCD to warming and expansion of shallow seasdue to a 50 m sea-level rise that led to increasedsequestration of CaCO3.

[26] There are also short (1–2 Ma) intervals at 19and 17 Ma when one oceanic mass significantlydiverges from another. These anomalous intervalstypically involve large negative excursions of Pa-cific benthic d13C, suggesting that there was inter-mittent production of a deep-water source, possiblylocated in the Antarctic or North Atlantic Oceans,depending on the direction of the d13C gradient.North Atlantic type surface water that sinks to formdeep water is enriched in 13C, but as deep watersmove away from their source region they graduallyaccumulate organic carbon rich in 12C from shal-

Figure 7. Analysis of effect of allowing for uncertain-ties in age determinations using 0–0.68 Ma portion ofPacific Ocean data. Solid circles, observed values; solidcurve, regression estimate; dashed curve, regressionestimate generated by SIMEX algorithm which removeseffect of uncertainty in observed age; gray band, 95%confidence interval for standard regression estimate.Effect of SIMEX procedure is to sharpen peaks andtroughs which are broadened and reduced in amplitudeif observations are poorly correlated due to datingerrors. Therefore age errors do not affect the generalshape of the curve (or of the resulting estimate of%NCW) but will affect the size of glacial-interglacialchanges in d13C and %NCW.



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Figure 8. Regression estimates for three sets of composite records. In each case, other two regression estimates areshown in pale shading for comparison. (a) North Atlantic Ocean. Solid black circles, observations; solid black curve,regression estimate; bands of different red shades, 95% confidence bands for each block of data. (b) Southern Ocean.As above with green shades. (c) Pacific Ocean. As above with different blue shades. Note overlap in confidencebands between 23 and 12 Ma.



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low waters. Thus its d13C signature is graduallyreduced as it travels from its source (i.e., water-mass aging).

[27] From 12–0 Ma, rapid divergence of the end-member curves may mark the beginning of moresustained large volume production of NCW. TheNorth Atlantic end-member oscillates around aconstant value from this time onward, reflectingcontinuous sinking of water depleted in 12C bysurface productivity. The Pacific Ocean end-member has the lowest values of d13C during thisperiod. The Southern Ocean end-member fluctu-ates between North Atlantic and Pacific Oceanvalues. For the purposes of this paper, we used13C as a water-mass tracer and assume thatchanges in Southern Ocean d13C are caused byvariations in the strength of NCW export from theNorwegian Sea. These variations can only give anindication of the relative flux of NCW in theSouthern Ocean and should not be regarded as ameasure of absolute export fluxes.

4.1. Estimation of Northern ComponentWater 0–23 Ma

[28] Following Wright and Miller [1996], we wishto use these records to estimate the variation inNCW overflow over the last 23 Ma. According toOppo and Fairbanks [1987], the percentage ofNCW present in the Southern Ocean is given byequation (1). This mixing equation is undoubtedlya simplistic way of representing the relationshipbetween these three water masses, but it can beregarded as a useful starting point. It can only beapplied when d13CSO � d13CPO and d13CNA >d13CPO. Thus before 12 Ma, when all three confi-dence bands overlap, we cannot calculate %NCW[contra Wright and Miller, 1996].

[29] Between 12 and 0 Ma, %NCW can be calcu-lated ( dNCW ) and an estimate of its variance isgiven by

bVdNCW ¼ A

d13CNA � d13CPO

� �2 þ B

d13CNA � d13CPO

� �4

� 2C

d13CNA � d13CPO

� �3 ð3Þ

where A = V SO + V PO, B = (d13CSO �d13CPO)

2(V NA + V PO) and C = (d13CSO �d13CPO)V

PO (see Serfling [1980] for furtherdetails). Here, V SO, V PO and V NA are theestimated variances of d13CSO, d

13CPO and d13CNA,respectively.

[30] %NCW is shown in Figures 9 and 10. Wemust emphasize that %NCW is unlikely to de-crease to zero at any time during the last 12 Maand so the percentage range is best thought of asindicating ‘‘weaker’’ and ‘‘stronger’’ NCW over-flow. When %NCW is close to zero, it may partlyindicate movement of the mixing zone northwardor shoaling of the mixing zone such that SouthernOcean sites no longer sit within this zone. Duringthe last 12 Ma, the most striking feature of calcu-lated %NCW is the rapid increase at 6 Ma anddecrease at 2 Ma. Other short period variations alsoappear to be well-resolved.

4.2. Evolutive Spectral Analysis

[31] The composite records span a large interval andwe expect that amplitudes of different frequencyresponses to insolation will vary with time. Theserecords have been carefully constructed and areof high sample resolution so it is useful to studychanges in power spectra of different orbital perio-dicities contained within the data. An evolutivespectral analysis technique is used to display tempo-ral changes in the power of different orbital perio-dicities. To apply this technique, results from thelocal linear kernel regression are exploited becausebandwidths were chosen to reflect sample resolution,enabling short-period cyclicities, where present, to beretained. In some age intervals, there is limitedspectral information for orbital frequencies ofinterest. Evolutive spectra are generated usingthe Thomson multi-taper method with a time-bandwidth product of 4 and eigenspectra of 6[Torrence and Compo, 1998]. Nonparametric regres-sion estimates were analyzed with a sliding windowlength of 1500 ka and a step size of 125 ka. Theshorter %NCW record was analyzed with a windowlength of 625 ka and a step size of 62.5 ka.

[32] Individual records show that, where there issignificant data resolution, nonparametric analysisof composite records retains the orbital cyclicitiesexpected from benthic stable isotope records. Wehave also examined the spectral content of the 0–7 Ma %NCW estimate since there are spectralpeaks in the benthic d13C data whence it is calcu-lated (Figure 10). The presence of obliquity isnotable between 3 and 1.6 Ma, although it ispresent quite consistently across the entire 0–7 Ma interval (Figure 11). This strong obliquitysignal suggests that %NCW is sensitive to obliq-uity forcing. Such a response may be characteristicof either the way carbon cycles respond to insola-tion forcing in the oceans during Late Neogene



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times, or a determinant of NCW formation in theNorthern Seas. It is interesting to note that theobliquity component of the North Atlantic end-member during this interval decays from a peakaround 6.5 Ma and is weak by 5 Ma. In contrast,the Pacific Ocean benthic d13C record only con-

tains a large obliquity component from �3.5 Ma.[Haug et al., 2005] suggest that changes in thesubarctic Pacific halocline, triggered by an obliq-uity minimum at �2.7 Ma, resulted in moisturerelease and in the Northern Hemisphere Glaciation(NHG). The %NCW record may be responding toobliquity forcing in the Pacific Ocean beginningaround 3 Ma. The timing of changes in obliquityamplitude in the Pacific benthic d13C recordoccur before the changes in the surface subarcticPacific Ocean. The increase in obliquity compo-nent in %NCW �3 Ma, however, does coincidewith Northern Hemisphere Glaciation. The Plio-Pleistocene benthic d18O stack of Lisiecki andRaymo [2005] also shows changes in orbitalforcing response, which occurred in conjunctionwith the climate transition �3 Ma. Their dataindicate a change in the phase (or lag) of theprecession and obliquity components of benthicd18O with respect to an orbitally forced ice-volumemodel at this time. Haug and Tiedemann [1998]suggest that NHG was caused by a combination ofthe late Cenozoic cooling trend, Panama Isthmusclosure affecting deep-water circulation in theNorth Atlantic, and finally triggered by favorableobliquity forcing. Recent modelling to test thishypothesis has failed to find support that thedelivery of moisture (through closure of the Pan-ama Isthmus) was a precondition for the obliquitytrigger [Klocker et al., 2005]. The timing of theimpact of closure of the Central American Seawayon deep-water circulation �4.6 Ma [Ravelo et al.,2004] also casts doubt on the closure of thePanama Isthmus as the final precondition forobliquity-triggered NHG [Haug et al., 2005].Klocker et al. [2005] suggest that whilst obliquitymodulation may be the trigger, the climate systemneeded different preconditions to allow the long-term buildup of ice in the Northern Hemisphere.

5. Iceland Plume Variation?

[33] Wright and Miller [1996] suggested that thevariation of NCW flux was controlled by verticalmotions of the GSR which were in turn controlledby temperature fluctuations within the IcelandPlume. Here we wish to test this hypothesis inthe light of our newly calculated %NCW and ourstudy of the associated errors (Figures 9 and 10).

5.1. Gateway Closure

[34] Given considerable uncertainties in paleogeo-graphical reconstructions, we are conscious that%NCW variation can be interpreted in various

Figure 9. (a) Calculated %NCW record for Neogeneperiod using data from Muller et al. [1991]. See section6.3 for a discussion of these data. Solid line, calculatedvariation in %NCW; gray shading, ±1s. (b) %NCWrecord with no Muller et al. [1991] data set; see Tables1–3. (c) Residual depth variation taken fromWright andMiller [1996] and shifted by +0.11 Ma to give bestnegative correlation to the %NCW record in Figure 9b.(d) Residual depth record shifted by +0.21 Ma torepresent largest shift possible to represent plumeanomalies positioned beneath Denmark Straits. Atime-varying correction might be more appropriate butcannot be fully justified at present. (e) Long wavelengthgravity anomalies from Irminger Basin shifted for theDenmark Straits. (f) Long-term variation of global d18Obased on local linear regression of oxygen data from alldata sets used here (Tables 1–3).



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ways. We are also conscious of recent suggestionsthat orbital forcing may combine with gradualchanges in atmospheric and circulation boundaryconditions to produce large, long-term changes inparameters such as benthic d18O and %NCW [e.g.,Haug and Tiedemann, 1998; Haug et al., 2005].Here, we restrict ourselves to tectonic explanations,although we acknowledge that there may be other(contributory) causes of %NCW variations.

[35] Two crucial issues affecting NCW productionare convective overturning within the Labrador Seaand changes in inter-oceanic communication. TheLabrador Sea is evidently an important source ofdeep water whose temporal evolution is poorlyunderstood. For now, we must assume that itscontribution does not vary significantly over theinterval of interest. On the other hand, opening andclosing of gateways could profoundly influencethermohaline circulation (including Labrador Seaconvective overturn) and we must address thegateway issue in more detail.

[36] Two significant gateway closures occur duringthe interval of interest: progressive closure of thePanama Gateway and final closure of the TethyanOcean. In both cases, deep-water circulation pat-terns and d13C gradients between the major oceanicbasins were undoubtedly affected. Tethyan oceanclosure is imprecisely dated but it probably influ-enced Atlantic deep waters until �13 Ma [Ramsayet al., 1998].

[37] Onset of significant NCW production mayhave been influenced by closure of the PanamaGateway, which gradually separated Atlantic andPacific oceanic circulation through Miocene-Pliocene times. Global circulation models have beenused to suggest that the present-day configurationof global thermohaline circulation is not possiblewith an open Panama gateway [Maier-Reimer et al.,1990]. If the pathway of warm, salty water associ-ated with the Gulf Stream is forced to the north, it ispossible that the Labrador Sea becomes the site ofimportant deep-water production. It is also likelythat, whilst deep-water production was possible in

Figure 10. Expanded portion of Figure 9 showing %NCW for 0–12 Ma interval. Black line, calculated %NCW;gray band, ±1s. Note that although calculated %NCW decreases close to zero, it is preferable to think of %NCWfluctuating from ‘‘stronger’’ to ‘‘weaker.’’



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the Norwegian–Greenland Seas prior to the for-mation of the Panama Isthmus, the flux wassignificantly increased by the Gulf Stream bringingwarm salty surface waters to high latitudes. Bothchanges could significantly influence the quantityof NCW exported into the Southern Ocean. Ourprincipal consideration here is timing of gatewayclosure. Foraminiferal stratigraphy of CentralAmerica suggests that initial uplift of the PanamaSill occurred at 15–15.9 Ma, although the deep-water connection was probably not severed at thistime [Duque-Caro, 1990]. Separation of Pacificand Caribbean benthic faunas, following a hiatusat 12.2–13.1 Ma, represents the first evidencefor development of a sill which was shallowenough to restrict Atlantic-Pacific deep water-mass exchange [Duque-Caro, 1990]. Nesbitt andYoung [1997] also argue that closure began atabout 12 Ma. Some deep water-mass exchangemay have continued until 6.7 Ma [McDougall,1996]. Studies of planktonic faunas indicate grad-ual separation of surface-water flow after thistime, suggesting complete divergence of surface-dwelling faunas by 1.8 Ma [Keller et al., 1989].Others have argued that water depths of lessthan 100 m were reached by 4.6 Ma on thebasis of surface salinity increases recorded by

planktonic foraminifera [Keigwin, 1982; Haugand Tiedemann, 1998].

[38] In summary, gradual restriction of deep-waterflow relatively early in the history of closure of thePanama gateway could have played a role inaffecting major NCW production during mid-lateMiocene times. Although there is some disagree-ment about the exact timings of shallowing andclosure, the consensus in the literature suggests thatshoaling to 100 m occurred at about 4.6 Ma [Haugand Tiedemann, 1998] and that complete closureoccurred at about 3.5 Ma [Duque-Caro, 1990;Haug and Tiedemann, 1998; Frank et al., 1999](Figure 10). Neither event appears to correspondwith major changes in the d13C composite recordsor %NCW.

5.2. Long-Term Subsidence of Greenland-Scotland Ridge

[39] Long-term subsidence of the GSR in responseto plate cooling could have exerted a primarycontrol on onset of NCW production [Vogt,1972]. Subsidence histories of the three present-day GSR gateways are considered separately sincethey have different tectonic histories (Figure 1).The Faroe-Shetland Channel is floored by a conti-

Figure 11. Thompson multitaper analysis of %NCW record for 0–7 Ma. Record shows persistent obliquity (i.e.,40 ka) influence, whereas eccentricity (100 ka) signal is patchy between 6.1 and 5.5 Ma but more persistent from2.8 Ma. Precessional signals are clearly evident between 1.8 and 0.8 Ma.



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nental extensional sedimentary basin. Drilling andseismic imaging by the hydrocarbon industryshow that this region subsided rapidly as a resultof Early-Middle Cretaceous rifting. It probablyremained at, or deeper than, its present depth of>1 km throughout the Cenozoic era, becoming aconduit for deep water by the early Oligocene[Davies et al., 2001]. It is therefore unlikely thatsubsidence of this gateway has significantly mod-erated production of NCW during the Neogeneperiod. The Iceland-Faroe Ridge (IFR) is built fromoverthickened oceanic crust overlain by a thinveneer of sediment 0–400 m thick [Smallwoodet al., 1999]. Its elevation at time tr before present,e, can be reconstructed by accounting for platecooling using

e ¼ h0 þ 380ffiffiffiffita

p�

ffiffiffiffiffiffiffiffiffiffiffiffiffita � tr

p� �ð4Þ

where ta is plate age in Ma, h0 is present elevationin meters, and constant 380 m Ma�1/2 is from theoceanic age-depth model [Stein and Stein, 1992].Using a maximum present-day water depth of500 m and a basement age range of 38–55 Ma[Smallwood and White, 2002], the IFR would havebeen elevated above sea level before 18–15 Ma, inagreement with Thiede and Eldholm [1983]. Thesubsidence history determined from DSDP hole336 on the northern flank of the ridge is compatiblewith this interpretation [Detrick et al., 1977]. By12 Ma, the IFR crest should have been 100–200 mbelow sea-level. The Denmark Straits are underlainby oceanic crust in the east and by the Greenlandpassive continental margin further west. Despite itscomplex basement architecture, subsidence ofDenmark Straits basement was probably similarto that of the IFR, since subsidence of 55 Maoceanic crust approximates postbreakup subsi-dence of the passive margin and oceanic crustbeneath the eastern part of the straits is about 38 Maold [Smallwood and White, 2002]. A larger sourceof uncertainty in sea-bed subsidence arises fromuncertainty in accumulation history of the 1 kmthick sedimentary layer which mantles DenmarkStraits basement [Wright and Miller, 1996]. Back-stripping this sediment layer shows that the seabedmay have been a maximum of 500 m deeper thanthe depth predicted by the simple plate coolingmodel above, if the entire sediment layer accumu-lated since the beginning of the Miocene period[Jones and Maclennan, 2005, Figure 1c]. It is morelikely that this sediment layer accumulated gradu-ally so that the seabed was never much more than100 m deeper than predicted by equation (4).

[40] We suggest that subsidence of the Greenland-Scotland Ridge on either side of Iceland exerted adirect control on NCW formation that was as leastas significant as closure of the Panama Gateway.At present, more than half of the northward flow ofsaline surface water travels across the Iceland-Faroe Ridge and more than half of the southwardreturn flow of NADW uses the Denmark Straits[Hansen et al., 2003]. The Iceland-Faroe Ridge,and possibly also the Denmark Straits, probably layabove sea-level prior to middle Miocene times.Limited connection between the Atlantic Oceanand the Greenland-Iceland-Norway Sea served toinhibit formation of NCW north of Iceland. Byabout 12 Ma, it is likely that water depths in excessof 100 m had been established above the GSRimmediately either side of Iceland, which wouldhave allowed saline Atlantic surface water to flownorthward, cool, and then sink to form NCW.Initially, return flow of NCW probably occurredthrough the Faroe-Shetland Channel, which hasprobably been open and relatively deep since EarlyOligocene times [Davies et al., 2001]. Ongoingsubsidence of oceanic crust around Iceland throughlate Miocene time allowed an increasing rate ofNCW export through the Denmark Straits, whichmay account for coeval divergence of the Atlanticand Pacific d13C records.

5.3. Short-Term Uplift/Subsidence ofGreenland-Scotland Ridge

[41] Short-lived cycles of uplift and subsidence aresuperimposed upon the long-term subsidence ofthe GSR. Wright and Miller [1996] used a set ofseismic reflection profiles that traverse the Rey-kjanes Ridge spreading axis to constrain the timingand amplitude of these cycles (Figures 1–3).Variations in depth to oceanic basement define aset of V-shaped ridges that are symmetric about thespreading axis and converge to the south [Vogt,1971]. A coeval set of V-shaped ridges that con-verge to the north occur along the KolbeinseyRidge, north of Iceland (Figure 1), [Jones et al.,2002a]. These V-shaped bathymetric ridges reflectcrustal thickness variations that are generated bythermal (and possibly compositional) anomalieswithin the head of the Iceland Plume that spreadoutward from Iceland. Fronts of relatively hightemperature, propagating outward from Iceland,are associated with dynamic uplift of the overlyingplate, since hotter regions are more buoyant. Thegeometry of the V-shaped ridges alone cannot beused to determine whether outward asthenosphereflow was radial or channelled beneath the spread-



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ing axis [White and Lovell, 1997; Jones et al.,2002a]. It is more likely that fronts of high tem-perature propagated radially beneath a horizontalboundary at a depth of about 100 km where a largeviscosity increase arises from dewatering of themantle in the earliest stages of melting [Ito et al.,1999; Ito, 2001; Jones et al., 2002a]. Hence V-shaped ridges formed at the Reykjanes Ridge southof Iceland can be correlated with uplift of the GSRto the west and east of Iceland.

[42] V-shaped ridge bathymetry can be convertedinto residual depth, which is the difference betweenthe observed depth to oceanic basement and thedepth predicted by a half-space cooling model.Depths to oceanic basement are derived fromTalwani et al. [1971], reported by Wright andMiller [1996]. Residual depth directly reflectsdynamic support, although features less than about0.1 Ma in duration may have been modified byfaulting and can be regarded as noise. Wright andMiller [1996] also presented estimates of accumu-lation rate of major North Atlantic sediment driftsand estimates of lava extrusion rates at Iceland assupporting evidence that vertical motion of theGSR affected the flow of NCW. Jones et al.[2002a] produced a V-shaped ridge record basedon gravity anomalies. The gravity record resemblesa smoother version of the residual depth record.

[43] The V-shaped records from residual depth andgravity were originally plotted as functions of timeusing magnetic anomalies correlated to the time-scales of Berggren et al. [1985] and Cande andKent [1995], respectively [Wright and Miller,1996; Jones et al., 2002a]. We have placed theserecords on the astronomical timescale of Lourens etal. [2004], since we wish to compare them with our%NCW record.

5.4. Timing of Uplift at the Major Spill-Points

[44] It is necessary to adjust the V-shaped ridgerecords to reflect the timing of uplift/subsidenceevents at locations of deep-water flow over theGSR in order to make a comparison with %NCW.Temperature anomalies that generate the V-shapedridges propagate radially outward from Iceland, sothe timing of associated uplift depends on thepropagation speed and radial distance from theplume center. The propagation speed is determinedfrom the angles between V-shaped ridges and thespreading axis. Wright and Miller [1996] measuredpropagation speeds between 1140 km/Ma for theyoungest ridge and 163 km/Ma for their oldest ridge.

[45] The residual depth record was formed at515 km from the plume center. Wright and Miller[1996] adjusted the original record to reflectuplift/subsidence at Iceland itself using a differentpropagation speed for each ridge. We removedthese adjustments to recreate the original record.We then sought a uniform time shift that resultedin a good negative correlation between rapidchanges in the slope of residual depth and%NCW. A uniform time shift of +0.11 Ma resultsin a reasonable negative correlation. DenmarkStraits and the Faroe Shetland Channel currentlylie 480 km and 760 km from the plume center,respectively. Using the range of propagationspeeds determined by Wright and Miller [1996],a time shift of between +0.03 and +0.21 Maapplied to the residual depth record should predictuplift/subsidence of the Denmark Straits, and ashift of between �0.22 and �1.52 Ma shouldpredict uplift/subsidence of the Faroe ShetlandChannel. The velocities found by Wright andMiller [1996] progressively increase for eachridge and so the time shift required for the plumeanomaly to reach Denmark Straits should be closerto +0.03 Ma for the younger ridges (Ridge 1) andcloser to 0.21 Ma for the older ridges (Ridges 3and 4). The residual depth record has been shiftedby +0.11 Ma (Figure 9c) and by +0.21 Ma(Figure 9d) to reflect vertical motions of theDenmark Straits. We have also tried time shiftswhich vary with age, as well as automated correla-tion. In our view, the use of an elastic correction isnot yet justified.

[46] Since a reasonable correlation between resid-ual depth and %NCW requires a uniform time shiftof +0.11, we suggest that vertical motions of theDenmark Straits have exerted the most importantcontrol on flow of deep water between the Norwe-gian-Greenland Sea and the Atlantic Ocean duringthe Neogene period. Comparing the V-shaped ridgegravity record with %NCW supports this conclu-sion, although the negative correlation is not sostriking because the gravity record is smoother thanthe residual depth record. Note that Wright andMiller [1996] and Jones et al. [2002a] found that%NCW was sensitive to uplift at Iceland itself. Theimproved quality of our %NCW record suggestsinstead that NCW flow is more sensitive to uplift atthe Denmark Straits.

6. Discussion

[47] Having made appropriate adjustments for thetime taken for a hot pulse to reach a distall spill



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point on the GSR (i.e., the Denmark Straits), wecan now compare our revised %NCW record withuplift/subsidence records of the Iceland Plume.Plume-related uplift and plate cooling correlatewith important changes in %NCW although weacknowledge that other events probably impact%NCW production and export during the Neogeneperiod.

6.1. %NCW Between 23 and 12 Ma

[48] Fluctuations in %NCW over the interval 23–12 Ma are inconclusive because the three compos-ite d13C records overlap within error. There isevidence for small inter-oceanic d13C gradientsbut carbon isotopes do not necessarily reveal thepresence of large-scale meridional circulation.Higher resolution data may help to reduce thewidth of confidence bands and to increase theprecision of correlation between end-members,but it is unlikely that large d13C gradients will beuncovered. The divergence of North Atlantic andPacific Ocean d13C records at 12 Ma is attributed tothe production of large volumes of NCW in theNorth Atlantic/Norwegian Greenland Sea. Deep-water formation patterns are more easily under-stood by examining statistically significant %NCWvariations.

6.2. Late Miocene to Present %NCWVariations: One Simple Cause?

[49] Since �12 Ma, large amounts of deep waterhave formed episodically in the North AtlanticOcean, causing d13C records for the North Atlanticand Pacific Oceans to diverge (Figure 8). SouthernOcean data fluctuate between these two end-members. Between 11 and 6 Ma, estimates of%NCW are reliable and negatively correlate withvertical motions of the GSR (Figure 9). This corre-lation is particularly striking when the residualdepth record is shifted to reflect uplift and subsi-dence of the Denmark Straits. Thus plume-relateduplift of Iceland appears to significantly restrict theamount of deep water which spills over the GSR atthe Denmark Straits. Its impact would be greaterthan simply reducing the 5.6 Sv of water whichcurrently overflow the ridge, since overflowedwater entrains intermediate waters which help toincrease the flux of present-day Lower NorthAtlantic Deep Water [Dickson and Brown, 1994].A weaker correlation between this %NCW recordand subsidence of the Faroe Shetland Channelsuggests that the Denmark Straits controlled large-volume NCW production.

[50] Figure 10 shows %NCW for the last 12 Mawhen the highest resolution records exist for eachof the three end-members. Several broad-scalefeatures are of interest. At 6 Ma, there was a rapidincrease in %NCW, which we interpret as anincrease in NCW production rather than a propor-tional increase in Southern Component Water.Another rapid increase occurred at 5.2 Ma. Thesechanges precede a sustained interval of NCWproduction which continues until 2.7 Ma. Weattribute the increase in %NCW at 6 Ma to reduceddynamic support at the Denmark Straits. Thisreduction is inferred from the trough betweenRidges 1 and 2. Wold [1994] has highlighted adramatic increase in sedimentary flux across theGSR at �6 Ma, and a slight decrease at �2 Ma,both of which correlate very well with an increaseand decrease in %NCW, respectively.

[51] From 2.7 Ma onward, the proportion of NCWin the Southern Ocean diminishes steadily until�1.9 Ma. Intensification of the Northern Hemi-sphere Glaciation and stagnation of waters in thenorthern seas also occurred at this time [Maslin etal., 1998]. This time interval coincides with theformation of V-shaped Ridge 1 in response to anincrease in temperature within the Iceland Plumehead. We suggest that the associated dynamic upliftof the GSR decreased southward overflow of deepwater from the NGS, and reduced strength ofthermohaline circulation. Stagnation of bottomwaters and poor carbonate preservation reportedby Henrich et al. [2002] ensued. Reduction inNCW also reduced latitudinal heat transport, there-by rapidly increasing the areal extent of NorthernHemisphere Glaciation.

[52] The composite benthic d18O record suggeststhat, despite large fluctuations in the proportion ofNCW in the Southern Ocean, periods of suddenincrease or decrease in %NCW are not generallyassociated with significant changes in deep-waterd18O composition. Thus long-term changes in thestrength of NCW do not obviously correlate withmajor changes in the Earth’s glaciation history.However, two of the major decreases in benthicd18O, at �15 and �3.6 Ma, do match the rapidgrowth of the two biggest V-shaped ridges formedduring the Neogene period (Ridges 1 and 3). Ourrecords suggest that changes in %NCW, causedby the formation of Ridge 1, are associated withan increase in Northern Hemisphere Glaciationat �3.6 Ma [Wright and Miller, 1996]. The possi-bility remains, given the uncertainties associatedwith %NCW prior to �12 Ma, that Ridge 3



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interrupted NCW production and export from theNorwegian-Greenland Seas, providing moisture forthe mid-Miocene growth in the Antarctic ice sheet[Schnitker, 1980].

6.3. Sensitivity of %NCW to Data Choice

[53] Although %NCW generally correlates wellwith Iceland Plume activity, there was initially asignificant mismatch at 7 Ma (Figure 9). Whencomposite records are scrutinized, it is clear thatthe 7 Ma peak in %NCW is caused by an increasein the Southern Ocean d13C composite record thatis controlled by a small number of high values(Figure 8). These high values of d13C come fromtwo sets of analyses carried out on ODP site 704 at2.5 km depth (Figure 12). These analyses divergefrom the trends in ODP site 1088 and DSDP site360 (at depths of 2 and 3 km, respectively). Thereis further disagreement between the two Site 704data sets, Muller et al. [1991] and Wright et al.[1991], where Muller’s data is generally positivelyoffset from the other data. Billups et al. [2002]states that there is no obvious environmental causefor the differences between the two Site 704 datasets that sample the late Miocene period. Wetherefore suggest that data from Site 1088 (andSite 360) better reflect the benthic d13C composi-tion of the Southern Ocean through this interval. IfSite 704 values are removed, causing the record tobe based on Site 1088 during the late Miocene, the%NCW peak at 7 Ma is considerably diminished(Figure 9). One cause of the difference in d13Cbetween Site 704 and Sites 360 and 1088 may besmall changes in the geometry of the mixing frontbetween NADW and CDW [Billups et al., 2002].The differences between Site 704 and Sites 360and 1088 highlight the importance of choosingappropriate data for each of the benthic d13C end-

members, since %NCW is sensitive to changes indata, especially in the Southern Ocean. Closeexamination of the records reveals no other regionof overlap between data sets with such a largediscrepancy.

7. Conclusions

[54] In order to test Wright and Miller’s [1996]assertion that temporal variations associated withthe Iceland Plume controlled overflow of NorthernComponent Water at the Greenland Iceland Ridge,we have assembled composite d13C records for theAtlantic, Southern and Pacific Oceans. These com-posite records were calibrated with the astronom-ical timescale and fitted with nonparametricfunctions. We paid particular attention to the con-struction of confidence bands.

[55] We have then used these three compositerecords to calculate the percentage overflow ofNorthern Component Water. There is no evidencefor significant NCW overflow before 12 Ma sincethe three oceanic end-members overlap withinerror. A marked divergence of the three d13Crecords at 12 Ma indicates the onset of significantNCW overflow which may be related to submer-gence of the Iceland-Faroe Ridge caused by litho-spheric plate cooling. This inference could be morerigorously tested by analyzing the detailed subsi-dence histories of the Greenland-Scotland Ridge,which requires deep drilling on the Iceland-FaroeRidge and in the Denmark Straits.

[56] Despite the many caveats concerning the useof d13C records to reconstruct overflow of NCW,and despite uncertainties surrounding the impor-tance of Labrador Sea convective overturn and thetiming of Panama Gateway closure, there is a

Figure 12. Detail of Southern Ocean d13C composite record which is highlighted in Figure 5b. This portion of datagenerates a peak in %NCW at about 7 Ma which does not agree with activity of Iceland Plume. Close inspectionreveal that ODP 704 data are shifting average values of d13C upward. If these data are removed, 7 Ma peak in %NCWdisappears.



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remarkably good correlation between %NCW andinferred vertical motions of the Denmark Straits.This correlation suggests that short-term (order 2–3 Ma) fluctuations of %NCW after 12 Ma arecontrolled by vertical motions of the GSR, whichare generated by temperature fluctuations associatedwith the Iceland Plume, vindicating the Wright andMiller hypothesis. Uplift fluctuations are likely tobe several hundred meters but detailed modelling isrequired to confirm this influence. The matchbetween short-term features of the %NCW recordand the V-shaped ridge record strengthens thesuggestion that the Greenland-Scotland Ridge alsocontrolled onset of NCW production in the longerterm. Nonetheless, more sustained production ofNCW may only occur after closure of the PanamaIsthmus in order to supply warm salty surface waterto high northern latitudes.

[57] Intensification of Northern Hemisphere Gla-ciation at 2.7 Ma may be related to the restrictionin %NCW in response to dynamic uplift associ-ated with the youngest V-shaped ridge. A reduc-tion in %NCW at 2.7 Ma may coincide with anincrease in obliquity forcing at 2.8 Ma. Thisreduction may indicate increased dominance ofNCW, which has a strong obliquity component, incontrolling circulation and glaciation. These hy-potheses could, in part, be tested by detailedexamination of North Atlantic ODP sites. Finally,the statistical techniques used to generate confi-dence bands for nonuniformly sampled isotopicdata may be of more general application inpaleoceanography.

Acknowledgments

[58] We warmly acknowledge the help of N. Shackleton,

K. Billups, and H. Pfuhl who very generously provided access

to unpublished data. L. Lourens, J. Ogg, H. Luterbacher, and

N. Shackleton gave us astronomical timescales, and H. Palike

helped construct evolutive spectral analyses. We also thank

S. Crowhurst, G. Edwards, C. Lear, D. Lyness, D. McKenzie,

and R. Rickaby for their considerable help and interest. The

thoughtful comments of two anonymous reviewers and the

editor, Laurent Labeyrie, greatly improved the manuscript.

Figures were produced using the GMT software package.

Department of Earth Sciences Contribution 8273.

References

Berggren, W. (1992), Neogene planktonic foraminifer magne-tostratigraphy of the southern Kerguelen Plateau (Sites 747,748 and 751), Proc. Ocean Drill. Program Sci. Results, 120,631–647.

Berggren, W., D. V. Kent, and J. A. van Couvering (1985), TheNeogene: Part 2, Neogene geochronology and chronostrati-

graphy, in The Chronology of the Geological Record, Geol.Soc. Mem., vol. 10, edited by N. Snelling, pp. 211–260,Blackwell Sci., Malden, Mass.

Billups, K. (2002), Late Miocene through early Pliocene deepwater circulation and climate change viewed from the sub-Antarctic South Atlantic, Palaeogeogr. Palaeoclimatol. Pa-laeoecol., 185, 287–307.

Billups, K., A. C. Ravelo, and J. C. Zachos (1997), EarlyPliocene deep-water circulation: Stable isotope evidencefor enhanced Northern Component Deep Water, Proc. OceanDrill. Program Sci. Results, 154, 319–330.

Billups, K., A. Ravelo, and J. Zachos (1998), Early Pliocenedeep water circulation in the western equatorial Atlantic:Implications for high-latitude climate change, Paleoceano-graphy, 13(1), 84–95.

Billups, K., J. E. T. Channell, and J. Zachos (2002), LateOligocene to early Miocene geochronology and paleoceano-graphy from the subantarctic South Atlantic, Paleoceanogra-phy, 17(1), 1004, doi:10.1029/2000PA000568.

Billups, K., H. Palike, J. Channell, J. Zachos, and N. Shackleton(2004), Astronomic calibration of the late Oligocene throughearly Miocene geomagnetic polarity time scale, Earth Planet.Sci. Lett., 224, 33–44.

Broecker, W., and T. Peng (1982), Tracers in the Sea, Lamont-Doherty Geol. Observ., Palisades, N. Y.

Burton, K., H. Ling, and R. O’Nions (1997), Closure of theCentral American Isthmus and its effect on deep-water for-mation in the North Atlantic, Nature, 386, 382–385.

Cande, S., and D. Kent (1995), Revised calibration of thegeomagnetic polarity timescale for the Late Cretaceous andCenozoic, J. Geophys. Res., 100(B4), 6093–6095.

Carroll, R., D. Ruppert, and L. Stefanski (1995), MeasurementError in Non-linear Models, CRC Press, Boca Raton, Fla.

Charles, C. D., and R. G. Fairbanks (1992), Evidence fromSouthern Ocean sediments for the effect of North Atlanticdeep-water flux on climate, Nature, 355, 416–419.

Claeskens, G., and I. van Keilegom (2003), Bootstrap confi-dence bands for regression curves and their derivatives, Ann.Stat., 31(6), 1852–1884.

Clement, B. M., and F. Robinson (1986), The magnetostrati-graphy of Leg 94 sediments, Initial Rep. Deep Sea Drill.Proj., 94, 635–650, Part 2.

Cook, J., and L. Stefanski (1994), Simulation-extrapolationestimation in parametric measurement error models, J. Am.Stat. Assoc., 89, 1314–1328.

Corfield, R. M., and J. E. Cartlidge (1993), Oxygen and carbonisotope stratigraphy of the middle Miocene, Holes 805B and806B, Proc. Ocean Drill. Program Sci. Results, 130, 307–322.

Curry, W., and K. Miller (1989), Oxygen and carbon isotopicvariation in Pliocene benthic foraminifers of the equatorialAtlantic, Proc. Ocean Drill. Program Sci. Results, 108,157–166.

Davies, R., J. Cartwright, J. Pike, and C. Line (2001), EarlyOligocene initiation of North Atlantic Deep Water formation,Nature, 410, 917–920.

Delaney, M., and E. Boyle (1987), Cd/Ca in late Miocenebenthic foraminifera and changes in the global organic car-bon budget, Nature, 330(6144), 156–159.

Detrick, R., J. Sclater, and J. Thiede (1977), The subsidence ofaseismic ridges, Earth Planet. Sci. Lett., 34, 185–196.

Dickson, R., and J. Brown (1994), The production of NorthAtlantic Deep Water: Sources, rates and pathways, J. Geo-phys. Res., 99(C6), 12,319–12,341.

Duque-Caro, H. (1990), Neogene stratigraphy, paleoceanogra-phy and paleobiogeography in northwest South America and



21 of 24

the evolution of the Panama Seaway, Palaeoceanogr. Pa-laeoclimatol. Palaeoecol., 77, 203–234.

Ellis, C. H. (1975), Calcareous nannofossil biostratigraphy -Leg 31, DSDP, Intial Rep. Deep Sea Drill. Proj., 31, 655–676.

Fan, J., and I. Gijbels (1996), Local Polynomial Modelling andIts Applications, chap. 2, pp. 16–19, CRC Press, Boca Ra-ton, Fla.

Frank, M., B. Reynolds, and R. O’Nions (1999), Nd and Pbisotopes in Atlantic and Pacific water masses before and afterclosure of the Panama gateway, Geology, 27(12), 1147–1150.

Gartner, S., and J. Chow (1985), Calcareous nannofossil bios-tratigraphy: Deep Sea Drilling Project Leg 85, eastern equa-torial Pacific, Initial Rep. Deep Sea Drill. Proj., 85, 609–619.

Graham, D. W., B. Corliss, M. Bender, and L. Keigwin (1981),Carbon and oxygen isotopic disequilibria of recent deep-seabenthic foraminifera, Mar. Micropaleontol., 6, 483–497.

Hailwood, E., and B. Clement (1991), Magnetostratigraphy ofSites 703 and 704, Meteor Rise, southeastern South Atlantic,Proc. Ocean Drill. Program Sci. Results, 114, 367–386.

Hansen, B., S. Østerhus, H. Hatun, R. Kristiansen, andM. Larsen (2003), The Iceland–Faroe inflow of Atlanticwater to the Nordic seas, Prog. Oceanogr., 59, 443–474.

Haug, G. H., and R. Tiedemann (1998), Effect of the formationof the Isthmus of Panama on Atlantic Ocean thermohalinecirculation, Nature, 393, 673–676.

Haug, G. H., et al. (2005), North Pacific seasonality and theglaciation of North America 2.7 million years ago, Nature,433, 821–825.

Heider, F., B. Leitner, and H. Inokuchi (1992), High southernlatitude magnetostratigraphy and rock magnetic properties ofsediments from Sites 747, 749 and 751, Proc. Ocean Drill.Program Sci. Results, 120, 225–245.

Henrich, R., K. Baumann, R. Huber, and H. Meggers (2002),Carbonate preservation records of the past 3 Myr in theNorwegian Greenland Sea and the northern North Atlantic:Implications for the history of NADW production, Mar.Geol., 184, 17–39.

Hodell, D., J. Curtis, F. Sierro, and M. Raymo (2001), Corre-lation of late Miocene to early Pliocene sequences betweenthe Mediterranean and North Atlantic, Paleoceanography,16, 155–163.

Hodell, D. A., and P. F. Ciesielski (1991), Stable isotopicand carbonate stratigraphy of the Late Pliocene andPleistocene of Hole 704A: Eastern subantarctic SouthAtlantic, Proc. Ocean Drill. Program Sci. Results, 114,409–435.

Imbrie, J., J. Hays, D. Martinson, A. McIntyre, A. Mix,J. Morley, N. Pisias, W. Prell, and N. Shackleton (1984),The orbital theory of Pleistocene climate: Support from arevised chronology of the marine d18O record, NATO ASISer., Ser. C, 126, 269–305.

Ito, G. (2001), Reykjanes ‘‘V’’-shaped ridges originating froma pulsing and dehydrating mantle plume, Nature, 411, 681–684.

Ito, G., Y. Shen, G. Hirth, and C. J. Wolfe (1999), Mantle flow,melting, and dehydration of the Iceland mantle plume, EarthPlanet. Sci. Lett., 165(1), 81–96.

Jenkins, D. G. (1978), Neogene planktonic foraminiferas fromDSDP Leg 40 Sites 360 and 362 in the southeastern Atlantic,Initial Rep. Deep Sea Drill. Proj., 40, 723–740.

Jones, S. M., and J. Maclennan (2005), Crustal flow beneathIceland, J. Geophys. Res., 110, B09410, doi:10.1029/2004JB003592.

Jones, S. M., N. White, and J. Maclennan (2002a), V-shapedridges around Iceland: Implications for spatial and temporalpatterns of mantle convection, Geochem. Geophys. Geosyst.,3(10), 1059, doi:10.1029/2002GC000361.

Jones, S. M., N. White, B. J. Clarke, E. Rowley, and K. Gallagher(2002b), Present and past influence of the Iceland Plume onsedimentation, in Exhumation of the North Atlantic Margin:Timing, Mechanisms and Implications for Petroleum Explora-tion, edited byA.Dore et al.,Geol. Soc. Spec. Publ.,196, 13–25.

Katz, M. E., D. R. Katz, J. D. Wright, K. G. Miller, D. K. Pak,N. J. Shackleton, and E. Thomas (2003), Early Cenozoicbenthic foraminiferal isotopes: Species reliability and inter-species correction factors, Paleoceanography, 18(2), 1024,doi:10.1029/2002PA000798.

Keigwin, L. (1982), Isotope paleoceanography of the Carib-bean and east Pacific: Role of Panama uplift in late Neogenetime, Science, 217, 350–353.

Keigwin, L., M.-P. Aubry, and D. Kent (1986), North Atlanticlate Miocene stable-isotope stratigraphy, biostratigraphy, andmagnetostratigraphy, Initial Rep. Deep Sea Drill. Proj., 94,935–963.

Keigwin, L. D. (1984), Stable isotope results on upper Mio-cene and Lower Pliocene foraminifers from Hole 552A, In-itial Rep. Deep Sea Drill. Proj., 81, 595–597.

Keller, G., C. Zenker, and S. Stone (1989), Late Neogenehistory of the Caribbean gateway, J. S. Am. Earth Sci., 2,73–108.

Klocker, A., M. Prange, and M. Schulz (2005), Testing theinfluence of the Central American Seaway on orbitallyforced Northern Hemisphere glaciation, Geophys. Res. Lett.,32, L03703, doi:10.1029/2004GL021564.

Kroopnick, P. (1985), The distribution of 13C of sCO2 in theworld oceans, Deep Sea Res., Part A, 32(1), 57–84.

Kuijpers, A., B. Hansen, V. Huhnerbach, B. Larsen, T. Nielsen,and F. Werner (2002), Norwegian Sea overflow through theFaroe–Shetland gateway as documented by its bedforms,Mar. Geol., 188(1–2), 147–164.

Lear, C. H., Y. Rosenthal, and J. Wright (2003), The closing ofa seaway: Ocean water masses and global climate change,Earth Planet. Sci. Lett., 210, 425–436.

Lisiecki, L. E., andM. E. Raymo (2005), APliocene-Pleistocenestack of 57 globally distributed benthic d18O records, Paleo-ceanography, 20, PA1003, doi:10.1029/2004PA001071.

Lourens, L., F. J. Hilgen, J. Laskar, N. Shackleton, andD. Wilson (2004), The Neogene Period, in A Geologic TimeScale 2004, chap. 21, pp. 409–440, Cambridge Univ. Press,New York.

Lyle, M. (2003), Neogene carbonate burial in the PacificOcean, Paleoceanography, 18(3), 1059, doi:10.1029/2002PA000777.

Mackensen, A., H.-W. Hubberten, T. Bickert, G. Fischer, andD. K. Ftterer (1993), The d13C in benthic for a foraminiferaltests of Fontbotia Wuellerstorfi (Schwager) relative to thed13C of dissolved inorganic carbon in Southern Ocean deepwater: Implications for glacial ocean circulation models, Pa-leoceanography, 8(5), 587–610.

Maier-Reimer, E., U. Mikolajewicz, and T. Crowley (1990),Ocean general circulation model sensitivity experiment withan open Central American Isthmus, Paleoceanography, 5,349–366.

Maslin, M., X. Li, M. Loutre, and A. Berger (1998), Thecontribution of orbital forcing to the progressive intensifica-tion of Northern Hemisphere Glaciation, Quat. Sci. Rev., 17,411–426.

McDougall, K. (1996), Benthic foraminiferal response to theemergence of the Isthmus of Panama and coincident pa-



22 of 24

leoceanographic changes, Mar. Micropaleontol., 28, 133–169.

Mikolajewicz, U., and T. Crowley (1997), Response of acoupled ocean/energy balance model to restricted flowthrough the Central American Isthmus, Paleoceanography,12, 429–441.

Miller, K., and B. Tucholke (1983), Development of Cenozoicabyssal circulation south of the Greenland-Scotland Ridge,in Structure and Development of the Greenland-ScotlandRidge, edited by M. Bott et al., pp. 549–589, Springer,New York.

Miller, K. G., and R. G. Fairbanks (1985), Oligocene to Mio-cene carbon isotope cycles and abyssal circulation change, inThe Carbon Cycle and Atmospheric CO2: Natural VariationsArchean to Present, Geophys. Monogr. Ser., vol. 32, editedby E. T. Sundquist and W. S. Broecker, pp. 469–486, AGU,Washington, D. C.

Miller, K. G., R. G. Fairbanks, and E. Thomas (1987), Benthicforaminiferal carbon isotopic records and the development ofabyssal circulation in the eastern North Atlantic, Initial Rep.Deep Sea Drill. Project, 94, 981–995.

Miller, K. G., M. D. Feigenson, J. D. Wright, and B. M.Clement (1991), Miocene isotope reference section, DeepSea Drilling Project Site 608: An evaluation of isotope andbiostratigraphic resolution, Paleoceanography, 6(1), 33–52.

Mix, A., and N. Shackleton (1995), Benthic foraminiferalstable isotope stratigraphy of Site 846: 0–1.8 Ma, Proc.Ocean Drill. Program Sci. Results, 138, 839–854.

Muller, D. W., D. Hodell, and P. Ciesielski (1991), Late Mio-cene to Earliest Pliocene (9.8–4.5 Ma) paleoceanography ofthe subantarctic southeast Atlantic: Stable isotopic, sedimen-tologic, and microfossil evidence, Proc. Ocean Drill. Pro-gram Sci. Results, 114, 459–474.

Nesbitt, H. W., and G. M. Young (1997), Sedimentation in theVenezuelan Basin, circulation in the Caribbean Sea, and on-set of Northern Hemisphere Glaciation, J. Geol., 105(531–544).

Oppo, D., M. Raymo, G. Lohmann, A. Mix, J. Wright, andW. Prell (1995), A d13C record of Upper North Atlantic DeepWater during the past 2.6 million years, Paleoceanography,11(3), 373–394.

Oppo, D. W., and R. G. Fairbanks (1987), Variability in thedeep and intermediate water circulation of the AtlanticOcean during the past 25,000 years: Northern hemispheremodulation of the Southern Ocean, Earth Planet. Sci. Lett.,86, 1–15.

Pfuhl, H., and N. Shackleton (2004), Two proximal, high-re-solution records of foraminiferal fragmentation and their im-plications for changes in dissolution, Deep Sea Res., Part I,51(6), 809–832, doi:10.10167j.dsr.2004.02.003.

Pfuhl, H. A., and N. I. McCave (2004), Integrated age modelsfor the early Oligocene-early Miocene, sites 1168 and 1170–1172, Proc. Ocean Drill. Program Sci. Results [CD-ROM],189, 22 pp.

Pisias, N., N. J. Shackleton, and M. A. Hall (1985), Stableisotope and calcium carbonate records from hydraulic pistoncored Hole 514A: High resolution records from the middleMiocene, Initial Rep. Deep Sea Drill. Proj., 85, 735–748.

Rahmstorf, S. (2002), Ocean circulation and climate during thepast 120,000 years, Nature, 419(6903), 207–214.

Ramsay, A. T., C. W. Smart, and J. C. Zachos (1998), A modelof early to middle Miocene deep ocean circulation for theAtlantic and Indian Oceans, in Geological Evolution ofOcean Basins: Results From the Ocean Drilling Program,edited by A. Cramp et al., Geol. Soc. Spec. Publ., 131, 55–70.

Ravelo, A., D. Andreason, M. Lyle, A. Olivarez Lyle, andM. Wara (2004), Regional climate shifts caused by gradualglobal cooling in the Pliocene epoch, Nature, 429, 263–267.

Raymo, M., D. Hodell, and E. Jansen (1992), Response ofdeep ocean circulation to initiation of Northern HemisphereGlaciation (3–2 Ma), Paleoceanography, 7(5), 645–672.

Ruddiman, W. (2001), Earth’s Climate: Past and Future, chap.11, pp. 235–253, W. H. Freeman, New York.

Ruddiman, W., M. E. Raymo, D. Martinson, B. Clement, andJ. Backman (1989), Pleistocene evolution: Northern hemi-sphere ice sheets and North Atlantic Ocean, Paleoceanogra-phy, 4(4), 353–412.

Samworth, R., and H. R. Poore (2005), Understanding pastocean circulations: A nonparametric regression case study,Stat. Modell., 5, 289–307.

Savin, S.M., R.Douglas,G.Keller, J.Killingley, L. Shaughnessy,M. Sommer, E. Vincent, and F. Woodruff (1981), Miocenebenthic foraminiferal isotope records: A synthesis, Mar.Micropaleontol., 6, 423–450.

Schmitz, W. J., Jr., and M. S. McCartney (1993), On the NorthAtlantic circulation, Rev. Geophys., 31(1), 29–49.

Schnitker, D. (1980), North Atlantic oceanography as possiblecause of Antarctic glaciation and eutrophication, Nature,284, 615–616.

Serfling, R. (1980), Approximation Theorems of MathematicalStatistics, chap. 3, p. 122, John Wiley, Hoboken, N. J.

Shackleton, N. (1987), The carbon isotope record of the Cen-ozoic: History of organic carbon burial and of oxygen in theocean and atmosphere, in Marine Petroleum Source Rocks,Geol. Soc. Spec. Publ., 26, 423–434.

Shackleton, N., and M. Hall (1984), Oxygen and carbon iso-tope stratigraphy of Deep Sea Drilling Project Hole 552A:Plio-Pleistocene glacial history, Initial Rep. Deep Sea Drill.Proj., 81, 599–610.

Shackleton, N., and N. Pisias (1985), Atmospheric carbondioxide, orbital forcing, and climate, in The Carbon Cycleand Atmospheric CO2: Natural Variations Archean to Pre-sent, Geophys. Monogr. Ser., vol. 32, edited by E. T.Sundquist and W. S. Broecker, pp. 303–317, AGU,Washington, D. C.

Shackleton, N., A. Berger, and W. Peltier (1990), An alterna-tive astronomical calibration of lower Pleistocene timescalebased on ODP Site 677, Trans. R. Soc. Edinburgh Earth Sci.,81, 251–261.

Shackleton, N., S. Crowhurst, T. Hagelberg, N. Pisias, andD. Schneider (1995a), A new Late Neogene Time Scale:Application to Leg 138 Sites, Proc. Ocean Drill. ProgramSci. Results, 138, 73–101.

Shackleton, N., M. Hall, and D. Pate (1995b), Pliocene stableisotope stratigraphy of Site 846, Proc. Ocean Drill. ProgramSci. Results, 138, 337–355.

Shackleton, N. J., and S. Crowhurst (1997), Sediment fluxesbased on an orbitally tuned time scale 5 Ma to 14 Ma, Site926, Proc. Ocean Drill. Program Sci. Results, 154, 69–82.

Shipboard Scientific Party (1975), Site 289, Initial Rep. DeepSea Drill. Proj., 30, 231–398.

Shipboard Scientific Party (1999), Site 1088, Proc. OceanDrill. Program Initial Rep. [CD-ROM], 177, 1–36.

Smallwood, J., and R. White (2002), Ridge-plume interactionin the North Atlantic and its influence on continental breakupand seafloor spreading, in The North Atlantic Igneous Pro-vince: Stratigraphy, Tectonic, Volcanic and Magmatic Pro-cesses, edited by D. Jolley and B. Bell, Geol. Soc. Spec.Publ., 197, pp. 15–37.

Smallwood, J. R., R. K. Staples, K. Richardson, and R. White(1999), Crust generated above the Iceland mantle plume:



23 of 24

From continental rift to oceanic spreading center, J. Geo-phys. Res., 104(B10), 22,885–22,902.

Srinivasan, M., and J. Kennett (1981), Neogene planktonicforamaniferal biostratigraphy and evolution: Equatorial tosubantarctic South Pacific, Mar. Micropaleontol., 6, 499–533.

Stein, C., and S. Stein (1992), Implications for ridges andtransforms of a hotter, thinner lithosphere, Eos Trans.AGU, 73, 550.

Takayama, T. (1993), Notes on Neogene calcareous nannofos-sil biostratigraphy of the Ontong Java Plateau and size var-iations of Reticulofenestra coccoliths, in Proc. Ocean Drill.Program Sci. Results, 130, 179–229.

Talwani, M., C. Windisch, and M. Langseth, Reykjanes Ridgecrest: A detailed geophysical study (1971), J. Geophys. Res.,76, 473–517.

Thiede, J., and O. Eldholm (1983), Speculation about the pa-leodepth of the Greenland-Scotland Ridge during Late Me-sozoic and Cenozoic times, in Structure and Development ofthe Greenland-Scotland Ridge, edited by M. Bott et al., pp.445–456, Springer, New York.

Tiedemann, R., and S. Franz (1997), Deep water circulation,chemistry and terrigenous sediment supply in the equator-ial Atlantic during the Pliocene, 3.3–2.6 ma and 5.4–5ma, Proc. Ocean Drill. Program Sci. Results, 154, 299–318.

Torrence, C., and G. P. Compo (1998), A practical guideto wavelet analysis, Bull. Am. Meteorol. Soc., 79(1),61–78.

van Andel, T. (1975), Mesozoic/Cenozoic calcite compensa-tion depth and the global distribution of calcareous sedi-ments, Earth Planet. Sci. Lett., 26, 187–194.

van Andel, T., et al. (1975), Cenozoic history and paleoceano-graphy of the central equatorial Pacific Ocean, inMem. Geol.Soc. Am., 143, 89–102.

Vincent, E., and W. Berger (1985), Carbon dioxide and polarcooling in the Miocene: The Monterey Hypothesis, in TheCarbon Cycle and Atmospheric CO2: Natural VariationsArchean to Present, Geophys. Monogr. Ser., vol. 32, editedby E. T. Sundquist and W. S. Broecker, pp. 455–468, AGU,Washington, D. C.

Vogt, P. (1971), Asthenosphere motion recorded by the oceanfloor south of Iceland, Earth Planet. Sci. Lett., 13, 153–160.

Vogt, P. (1972), The Faeroe-Iceland-Greenland aseismic ridgeand the western boundary undercurrent, Nature, 239, 79–81.

White, N., and B. Lovell (1997), Measuring the pulse of aplume with the sedimentary record, Nature, 387, 888–891.

Whitman, J. M., and W. H. Berger (1993), Pliocene-Pleisto-cene carbon isotope record, Site 586, Ontong Java Plateau, inProc. Ocean Drill. Program Sci. Results, 130, 333–348.

Wold, C. (1994), Cenozoic sediment accumulation on drifts inthe northern North Atlantic, Paleoceanography, 9, 917–941.

Woodruff, F., and S. M. Savin (1989), Miocene deepwateroceanography, Paleoceanography, 4(10), 87–140.

Woodruff, F., and S. M. Savin (1991), Mid-Miocene isotopestratigraphy in the deep sea: High-resolution correlations,paleoclimatic cycles, and sediment preservation, Paleocea-nography, 6(6), 755–806.

Wright, J. D., and K. G. Miller (1992), Miocene stable isotopestratigraphy, Site 747, Kerguelen Plateau, Proc. Ocean Drill.Program Sci. Results, 120, 855–864.

Wright, J. D., and K. G. Miller (1993), Southern Ocean influ-ences on late Eocene to Miocene deepwater circulation, inThe Antarctic Paleoenvironment: A Perspective on GlobalChange, Part Two, Antarct. Res. Ser., vol. 60, edited by J. P.Kenneth and D. A. Warnke, pp. 1–25, AGU, Washington, D.C.

Wright, J. D., and K. G. Miller (1996), Control of NorthAtlantic Deep Water circulation by the Greenland-ScotlandRidge, Paleoceanography, 11(2), 157–170.

Wright, J. D., K. G. Miller, and R. G. Fairbanks (1991), Evo-lution of modern deepwater circulation: Evidence from thelate Miocene Southern Ocean, Paleoceanography, 6(2),275–290.

Wright, J. D., et al. (1992), Early and middle Miocene stableisotopes: Implications for deepwater circulation and climate,Paleoceanography, 7(3), 357–389.

Xia, Y. (1998), Bias-corrected confidence bands in nonpara-metric regression, J. R. Stat. Soc., Ser. B, 60, 797–811.

Zachos, J. C., N. Shackleton, J. Revenaugh, H. Palike, andB. Flower (2001), Climate response to orbital forcing acrossthe Oligocene-Miocene boundary, Science, 292, 274–278.



24 of 24

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