Coccolithophore carbonate during the last 450 ka inthe NW Pacific Ocean (ODP site 1209B, Shatsky Rise)

MANUELA BORDIGA,1,2 MIRIAM COBIANCHI,1 CLAUDIA LUPI,1* NICOLA PELOSI,3

NICHOLAS L. VENTI4 and PATRIZIA ZIVERI5,61Dipartimento di Scienze della Terra e dell’Ambiente, Universita degli Studi di Pavia, Pavia, Italy2Department of Earth Sciences, Uppsala Universitet, Uppsala, Sweden3Istituto Ambiente Marino Costiero (IAMC) – Calata Porta di Massa (Interno Porto di Napoli, Napoli), Italy4University of Massachusetts-Amherst, Department of Geoscience, University of Massachusetts Amherst, Amherst, MA, USA5Universitat Autonoma de Barcelona, Insitute of Environmental Science and Technology, Edifici Cn Campus de la UAB,Bellaterra, Barcelona, Spain

6Earth & Climate Cluster, Department of Earth Sciences, Faculty of Earth and Life Sciences, VU Universiteit Amsterdam, deBoelelaan 1085, 1081HV Amsterdam, The Netherlands

Received 7 February 2013; Revised 21 November 2013; Accepted 24 November 2013

ABSTRACT: This study focuses on Pleistocene–Holocene sediments from the Shatsky Rise (Ocean Drilling ProgramSite 1209B, NW Pacific Ocean). We quantify the contribution of calcite made by coccoliths in oceanic sediments,investigating the role of calcareous nannoplankton during the last 450 ka. Coccolith carbonate constitutes 60–90% ofbulk carbonate. Coccolith carbonate accumulation rates (CARs) and CaCO3 fine fraction confirm that coccolithophoresare major contributors to the carbonate export and accumulation. Primary productivity shows highest values frommarine isotope stage (MIS) 12 to 8. Thereafter, although coccolith calcite content remains high, other sedimentcomponents, such as as foraminifera and biogenic opal, seem to be favored, perhaps related to an increase infertilization by eolian dust. Our results demonstrate the important role of coccolithophore production and sedimentationon the regulation of ocean carbonate chemistry on time scales >1000–100 000 years. On glacial–interglacial scales,coccolithophore productivity could have affected deep-water saturation by buffering deep-sea CO2 through increasedcarbonate dissolution episodes. Spectral and wavelet analyses are consistent with CARs primarily driven by glacial–interglacial variability and obliquity-controlled changes. Coccolith-based paleoceanographic reconstructions allow us toestablish that during the last 450 ka the mid-latitudes of the NW Pacific are controlled by the dynamics of the El Ni~noSouthern Oscillation perturbations and Boreal Monsoon system. Copyright # 2013 John Wiley & Sons, Ltd.

KEYWORDS: calcareous nannofossils; coccolith carbonate accumulation rate; NW Pacific Ocean; Pleistocene; Shatsky Rise.

Introduction

The oceans regulate the atmospheric CO2 concentrationsthrough surface water calcium carbonate production andexport, and deep-sea sediment accumulation (Ridgwell andZeebe, 2005; Palike et al., 2012). This mechanism is impor-tant at time scales >1–100 ka (Zeebe, 2012). As calcificationin the oceans is mainly controlled by marine calcifying biota(e.g. Martin, 1990; Sarmiento and Le Quere, 1996; Boyd etal., 2000), studying the contribution of different calcifyingfunctional groups provides new insights into their roles andfeedbacks to the carbon cycle. Among marine biota, calcare-ous nannoplankton have been major producers of pelagiccarbonates since the Mesozoic (e.g. Westbroek et al., 1993;Bown, 1998; Ridgwell and Zeebe, 2005).Coccolithophores are not only a major component of

carbonate export to the deep sea, but also perform photosyn-thesis contributing to the biological pump. They can alsoenhance carbon export by providing carbonate particlesacting as ballasting minerals (Klaas and Archer, 2003; Ziveriet al., 2007; Wilson et al., 2013). Several studies on modernand past sediments have pointed out that coccolithophoreCaCO3 production can account for 20–80% of biogeniccarbonate exported from shallower waters to deep-sea sedi-ments (e.g. Bornemann et al., 2003, 2004; Hay, 2004; Ziveriet al., 2007; Broecker and Clark, 2009).To trace the past role of coccolithophores in oceanic

carbonate export and sedimentation, it is important to quantifytheir contribution to total CaCO3, taking into account export

production, dissolution and sedimentary dilution processes thatcan influence the total CaCO3 content. Here we explore the last450 ka of the ODP core 198-1209B, located in the NW PacificOcean on Shatsky Rise, within a transitional zone betweensubtropical and subarctic waters (Fig. 1). This time interval andoceanographic setting are ideal to investigate Late Quaternarycarbonate export and sedimentation because the studied recordencompasses the Mid-Brunhes Event (MBE; Jansen et al., 1986),a period of significant perturbation to the global ocean carbonsystem; the transitional setting facilitates investigation of oceanicfront dynamics and related oceanographic features (Sancettaand Silvestri, 1986; Haug et al., 1995).In the Shatsky Rise area, modern seasonality of biogenic

particle fluxes has been the subject of many investigations(e.g. Noriki and Tsunogai, 1986; Kawahata et al., 2000;Kuroyanagi et al., 2002; Eguchi et al., 2003; Mohiuddinet al., 2002, 2004). Nonetheless, a few studies describe LateQuaternary fluctuations in biogenic, organic matter, andplanktonic and benthic foraminiferal content (Ujiie, 2003;Yamane, 2003), and only one has reconstructed the surfacewater conditions in the NW Pacific using calcareous nanno-fossil data (Chiyonobu et al., 2012). Some authors underlinea gap between the carbonate content in bulk (up to 90% byweight) and foraminiferal contribution (�30% by weight;Mohiuddin et al., 2002). Therefore, detailed studies on therole of nannofossils in carbonate export could be useful.This paper focuses on the carbonate dynamics of the last

450 ka in the NW Pacific Ocean, to (i) evaluate thecontribution of calcareous nannoplankton to carbonate sedi-ments in this area and (ii) reconstruct Pacific subtropical andsubarctic front evolution and ventilation changes influencingcarbonate dynamics during this time interval.

�Correspondence: C. Lupi, as above.E-mail: claudia.lupi@unipv.it

Copyright # 2013 John Wiley & Sons, Ltd.

JOURNAL OF QUATERNARY SCIENCE (2014) 29(1) 57–69 ISSN 0267-8179. DOI: 10.1002/jqs.2677

Oceanographic setting

ODP Site 1209B (32˚39.10810N, 158˚30.35640E) was recov-ered in the NW Pacific Ocean on Shatsky Rise, a plateau eastof Japan (Fig. 1). The site was cored at a middle bathyal waterdepth (2387m), above the regional lysocline (2900m;Vincent, 1975) and calcite compensation depth (4000–4500m; Berger and Winterer, 1974).Buffered by the Kuroshio Current Extension’s warmth, the

oligotrophic subtropical water mass characterizes surfaceconditions at this site throughout most of the year in themodern ocean (Mizuno and White, 1983; Qiu, 2001; Anto-nov et al., 2010). In winter, however, cold westerly winds,enhanced by the East Asian monsoon (Kalnay et al., 1996),prodigiously advect sensible (Da Silva et al., 1994) and latent(Xie and Arkin, 1997) heat from the Kuroshio Extension to theatmosphere, thereby cooling the surface ocean flow tovirtually erase the thermocline (Levitus and Boyer, 1994;Antonov et al., 2010). Resulting weakened stratification of thewater column is characteristic of the transitional water mass,in which nutrients can exchange more freely with surfacewaters, affording increased chlorophyll concentrations (e.g.Chai et al., 2003) and overall productivity when overheadinsolation increases in spring (Mohiuddin et al., 2002, 2004;Eguchi et al., 2003; Harada et al., 2006).At 300–600m water depth, below the Kuroshio Current,

the North Pacific Intermediate Water originates along thePolar Front. Finally, below 2000m, Pacific waters attain theuniform properties of Pacific Deep Waters (PDW), a mixtureof both Antarctic Bottom Water and North Atlantic DeepWater (Millero, 1996). Because access of Arctic BottomWater is prevented by the shallow Bering Strait(Warren, 1983; Keigwin, 1987; Broecker, 1995), PDW flowsinto the deep basins of the Pacific Ocean from the south-westvia the Deep Western Boundary Current.

Materials and methods

Age model

Fifty-nine samples of the uppermost 6m below the sea floor(mbsf) at Site 1209B were studied with a sampling interval of10 cm. In this section, sediment lithology mainly consists of

nannofossil oozes with a smaller, variable clay component(Bralower et al., 2002).The age model (Bordiga et al., 2013) is a combination of

19 control points of the oxygen isotope records (planktonicforaminiferal species Globoconella inflata and endobenthicspecies Uvigerina peregrina; Fig. 2; Table 1), tuned with theLR04 curve of Lisiecki and Raymo (2005), and two biostrati-graphic events [first occurrence of Emiliania huxleyi and thecross-over of Gephyrocapsa caribbeanica and E. huxleyi].Table 1 shows the age control points used in the calibration:depth (mbsf) follows Bordiga et al. (2013) and standard agesof bioevents are from Thierstein et al. (1977).Sample age between control points was calculated by

linear interpolation. The age model (Fig. 2) shows a continu-ous sedimentary succession which extends from marineisotope stage (MIS) 1 to 12 at 450 ka. Average sedimentationrate is 1.36 cm ka�1, and mean sampling resolution is7.36 ka.

Nannofossil assemblages, nannofossilaccumulation rate and coccolith-carbonatecontribution

To evaluate coccolith carbonate fluxes, we here prepared the59 samples using the filtration technique (Ziveri et al., 1999;Incarbona et al., 2011), as opposed to the settling techniqueemployed by Bordiga et al. (2013). The samples wereanalysed by counting nannofossil assemblages and evaluationof absolute coccolith concentrations (coccoliths g�1). Usingpolarized light microscopy (1250�), all specimens in tenfields of view were counted (cf. Incarbona et al., 2011). Thetotal number of specimens counted ranged between 200 and2600, depending on coccolith abundance. Specimens wereidentified to species level following calcareous nannofossilstaxonomies of Young (1998), Hine and Weaver (1998),Young et al. (2003) and references within //nannotax.orgedited by Young et al. (2011).We calculated coccolith relative abundance and concen-

tration (Cc) following the formula given by Incarbona et al.(2011), expressing these in percentage and as the number ofcoccoliths per gram of dry bulk sediment (N g�1; Lototskayaet al., 1998).Nannofossil accumulation rate (NAR; N cm�2 ka�1) was

calculated as in Ziveri et al. (1999) using: number ofcoccoliths per gram (Cc; N g�1), sedimentation rates (cmka�1) derived from the age model, and dry bulk density(g cm�3) obtained by linear interpolation between dry bulkdensity data measured using a hydrostatic balance on sixsamples (with a spacing of 1m).The coccolith carbonate mass (Ncarb) related to each species’

abundance was calculated by a cubic function as presented inYoung and Ziveri (2000) and Bornemann et al. (2003):

Ncarb ¼ Cc� r� l3 � ks

where Cc is coccolith concentration (N g�1), r is calcitedensity (2.7 g cm�3); l is species-specific mean coccolithlength and ks is a species-specific shape factor.Table 2 summarizes species-specific shape factor (ks) and

mean coccolith length values (l) used in this work. For everyspecies at least 20 specimens were measured by lightmicroscopy to obtain an average length. This method allowsestimation of coccolith-derived CaCO3 from individual spe-cies, although significant cumulative error (around 50%) isinvolved (Young and Ziveri, 2000). For several samples, onepotential source of error is variation in size of specimens

Figure 1. Locations of study Site ODP 1209B and nearby sites cited inthe text. Also shown are the current trajectory of the KuroshioExtension and Oyashio Current and the position of the Subtropical andSubarctic Fronts (modified from Mizuno and White, 1983; Kawahataand Ohshima, 2002).

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within the same species (Sprengel et al., 2000). Moreover,broken coccoliths are not taken into account, and hencecoccolith carbonate contribution could be higher than com-puted (Lo�pez-Otalvaro et al., 2008). Thus, calculation ofcoccolith carbonate content should be considered an under-estimation of the total coccolith carbonate contribution(Young and Ziveri, 2000; Ziveri et al., 2007).Finally, coccolith-carbonate flux (coccolith accumulation

rate, CAR) evaluated from light microscopy counts (CARLM; gcm�2 ka�1) was calculated by relating coccolith-CaCO3 mass(Ncarb) to sedimentation rate and dry-bulk density, as pre-

sented in Young and Ziveri (2000) and Bornemann et al.(2003).

Sr/Ca analysis and carbonate content (wt%)

One gram of dry-bulk sediment was sieved through a <20-mmsieve and the fine fraction weighed to evaluate the amount ofthis fraction, i.e. the calcareous nannofossils (Chiu andBroecker, 2008), relative to bulk sediment composition. Sr/Caratio was measured in this fine fraction via inductively coupledplasma-mass spectrometry in pulse-counting mode on thediluted solution at the Department of Analytical Chemistry(Oviedo University). Matrix-matched standards were used tocorrect for mass bias and detector dead time. The instrumentalprecision for Sr/Ca coccolith has a reproducibility better than0.02mmolmol�1. The ratio was calibrated following Stoll et al.(2002), employing sea surface temperature (SST) estimates fromODP Site 1208, some 4˚ north of Site 1209 in the KuroshioExtension (Fig. 1; LaRiviere et al., 2012). Since Sr/Ca valueswere measured on the <20-mm fraction, they provide apotential coccolith productivity proxy as several authors havehighlighted (e.g. Stoll and Schrag, 2000, 2001; Stoll et al.,2002; Barker et al., 2006; Rickaby et al., 2007).In addition, Sr/Ca values were used to calculate CaCO3

(wt%) on the <20-mm fraction, using the Ca3158 valuesproportionated to the initial bulk-weight (400–500mg) em-ployed for the geochemical analysis.CaCO3 content of the bulk sediment was measured in all

the samples by gasometric methods (Muller andGastner, 1971) and expressed in percentage by weight (wt%).CaCO3wt% of both the bulk and <20-mm fraction wereemployed to calculate CAR (g cm�2 ka�1) for both bulk(CARbulk), and the <20-mm fraction (CAR<20mm) as follows:

CAR ¼ ðDBD� SR� CaCO3Þ=100

where DBD is the dry-bulk density (g cm�3) and SR issedimentation rate (cm ka�1).

Figure 2. Age model for ODP Site 1209B (Bordiga et al., 2013). Benthic and planktic foraminifer oxygen isotope record of the ODP Site 1209Btuned to the orbitally calibrated LR04 reference stack (Lisiecki and Raymo, 2005). Biostratigraphic events: E. huxleyi First Occurrence (FO) and G.caribbeanica-E. huxleyi cross-over. Partial and linear sedimentation rates and tie points are reported: black circles and crosses identify the tie pointsderived from the benthic and planktic d18O curves, respectively (Table 1). Numbers around d18O curve indicate marine oxygen isotope stages (MIS);grey bars show glacial intervals.

Table 1. Age control points for ODP Site 1209B.

Depth (mbsf) Age (ka) Data source

0.1 9 A0.2 18 A; C0.5 73 A0.7 83 A1 109 A1.2 125 A1.7 143 A2 182 A2.2 198 A2.5 224 A2.6 239 A3 254 A3.2 267 A; D3.5 283 A3.7 299 B4.7 331 B4.8 342 B5.3 411 B5.8 439 B

A, benthic foraminifera d18O event; B, planktic foraminifera d18Oevent; C, cross-over of Gephyrocapsa caribbeanica and Emilianiahuxleyi; D, first appearance datum of E. huxleyi. Reference ages ofthe bioevents are from Thierstein et al. (1977).

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COCCOLITHOPHORE CARBONATE IN THE NW PACIFIC OCEAN 59

Paleoproductivity and carbonate preservationindices

Coccolith species relative abundance was used to calculate apaleoproductivity proxy, i.e. the N ratio of Flores et al. (2000)in place of the Florisphaera profunda index used in Bordigaet al. (2013). The N ratio is based on the relative abundanceof the main surface eutrophic species (small Noelerhabda-ceae) and F. profunda. High values imply a shallowernutricline/thermocline and higher surface productivity; con-versely, low values indicate a relatively deep nutricline/thermocline and deep stratification of the water column.The nannofossil dissolution index (NDI), provided by

Bordiga et al. (2013), is applied to evaluate carbonatepreservation. NDI (Marino et al., 2009) considers the ratiobetween proximal and distal shield abundances of Calcidis-cus with respect to the entire placoliths of the taxon, and itwas calculated by counting all subspecies of Calcidiscusleptoporus. This index shows high values for excellentpreservation and has been demonstrated to be a reliableand reproducible nannofossil dissolution proxy (Bordigaet al., 2013).

Spectral analysis

Spectral analysis was employed to identify significant perio-dicities and their evolution in the ODP Site 1209B total NARand CARLM time series. To accommodate uneven sampling,REDFIT software (Schulz and Mudelsee, 2002) and Foster’swavelet analysis algorithm were applied (Foster, 1996a,1996b, c). REDFIT software allows a test of significance inspectrum peaks against the null hypothesis of a red-noisebackground, which is estimated using a first-order autoregres-sive signal with sampling times and characteristic time-scalesmatching those of the real data without requiring interpola-tion. Wavelet analysis is useful for analysis of non-stationarysignals: frequency changes over time, or time-dependentmeasurement of amplitude and period of dominant cycles. Itcircumvents the non-stationary problem by expanding a timesignal in terms of a set of basic mathematical functionslocalized in both time and frequency domains.

Results

Calcareous nannofossil assemblages

The total coccolith concentration (total Cc), expressed as thenumber of coccoliths per gram of dry sediment, ranges fromabout 0.5 to 4� 1010 with a maximum at the MIS 10/MIS 9transition, and minimum values within MIS 8, MIS 7 and MIS5 (Fig. 3). Total Cc fluctuates on glacial–interglacial cycles,displaying higher values mainly during glacial intervals.Before MIS 8, assemblages are dominated by small

Gephyrocapsa caribbeanica approching abundance of 80%.Our record, in fact, falls within the G. caribbeanica acme(560–280 ka; Baumann and Freitag, 2004). Thereafter, thecoccolith assemblage is dominated by G. aperta, G. oceanicaand F. profunda, collectively comprising >80% of theassemblage. Other well-represented taxa include intermediateCalcidiscus leptoporus (abundance up to 20%) and G.muellerae (up to 16%). Small Calcidiscus leptoporus (up to8%) increases sharply during MIS 7. Finally, C. quadriperfor-atus represents only 4% of the assemblages. Distinct subspe-cies of Coccolithus pelagicus have been recently recognized(Geisen et al., 2002; Jordan et al., 2004; Parente et al., 2004).Most abundant here is C. braarudii (up to 6%), followed byC. pelagicus pelagicus (up to 5%) and by very rare specimensof C. pelagicus azorinus (up to 1.6%). The relative abundanceof Helicosphaera carteri increases up-section, reaching maxi-mum values close to 8% at the end of MIS 6 and during MIS5. Finally, the assemblages contain other taxa, such asRhabdosphaera clavigera, Syracosphaera spp., Umbellos-phaera spp. and Umbilicosphaera sibogae (not shown inFig. 3). Aizawa et al. (2004) grouped these taxa together,providing a ‘Warm Water Taxa index’. However, in oursamples, they are discontinuous and all have abundances<4%, preventing their use as a meaningful tool to examineSST change.

Paleoproductivity and NARs

The N ratio record exhibits high values throughout thestudied interval, except for MIS 7; in particular, highest values

Table 2. Summary of coccolith shape constants (ks) and mean lengths employed herein.

Species ks ks source Medium length (mm)

Calcidiscus leptoporus 0.08 Young and Ziveri (2000) 7Calcidiscus leptoporus small 0.105 Young and Ziveri (2000) 4.3Calcidiscus quadriperforatus 0.069 Young and Ziveri (2000) 9Coccolithus pelagicus azorinus 0.06 Young and Ziveri (2000) 15Coccolithus pelagicus braarudii 0.06 Young and Ziveri (2000) 13Coccolithus pelagicus pelagicus 0.06 Young and Ziveri (2000) 8.5Emiliania huxleyi 0.02 Young and Ziveri (2000) 3Florisphaera profunda 0.03 Young and Ziveri (2000) 3.3Medium G. caribbeanica 0.06 Lo�pez-Otalvaro et al. (2008) 4.5Medium G. oceanica 0.05 Young and Ziveri (2000) 4.5Gephyrocapsa aperta 0.035 Adopted from small placoliths (Young and Ziveri, 2000) 2Gephyrocapsa caribbeanica small 0.06 Lo�pez-Otalvaro et al. (2008) 2.8Gephyrocapsa ericsonii 0.05 Young and Ziveri (2000) 2Gephyrocapsa muellerae 0.05 Young and Ziveri (2000) 3.7Gephyrocapsa oceanica small 0.05 Young and Ziveri (2000) 3Small Gephyrocapsa spp. 0.035 Young and Ziveri (2000) 2.4Helicosphaera carteri 0.05 Young and Ziveri (2000) 9.8Pontosphaera spp. 0.05 Young and Ziveri (2000) 9.5Rhabdosphaera clavigera 0.025 Young and Ziveri (2000) 8Syracosphaera histrica 0.03 Young and Ziveri (2000) 5Syracosphaera pulchra 0.03 Young and Ziveri (2000) 5Umbilicosphaera foliosa 0.06 Young and Ziveri (2000) 4.5Umbilicosphaera sibogae 0.055 Young and Ziveri (2000) 4.5

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are recorded from the bottom of the study interval to MIS 8(Fig. 4). The Sr/Ca ratio, documenting the role of calcareousnannofossils in the biological pump (Stoll and Schrag,2000, 2001; Ziveri et al., 2000; Barker et al., 2006; Rickabyet al., 2007), shows ample fluctuation on the glacial–intergla-cial scale, with maximum values during interglacials. Twodistinct minima fall within MIS 8 and at the MIS 5/MIS 4transition. Total NAR ranges from 1� 1010 to 8�1010

nannofossils cm�2 ka�1 and shows a decreasing long-termtrend up-section. At the glacial–interglacial scale, total NARrecords positive peaks mainly at the end of glacials andduring interglacial intervals.From MIS 12 to 8, the chief contribution to total NAR is

from G. caribbeanica, which experiences its acme during thisinterval, corresponding to the MBE (Jansen et al., 1986).After the decline of G. caribbeanica, total NAR reaches

low values with episodic peaks, particularly during deglacia-tions at the MIS 8/MIS 7, MIS 6/MIS 5 and MIS 2/MIS 1transitions. Chief contributors to the oldest episode (MIS 8/MIS 7) are C. leptoporus, F. profunda, G. muellerae, H.carteri and the different subspecies of C. pelagicus. The totalNAR peak at the MIS 6/MIS 5 transition is mainly due to G.aperta, G. muellerae, C. leptoporus intermediate, G. oceanicaand H. carteri. Finally, the younger peak (MIS 2/MIS 1) isrelated to high concentrations of F. profunda, H. carteri, G.oceanica, C. pelagicus azorinus and the different subspeciesof C. leptoporus.

Carbonate (coccolith) and non-carbonate (dilution)fluxes and preservation

The carbonate content of deep-sea sediments is the result ofinterplaying processes: export–production of CaCO3-secretingorganisms in surface waters, CaCO3 solubility in seawatersand at the seawater/sediment interface, and dilution by non-carbonate or terrigenous materials.Figure 5 shows percentage (by weight) of carbonate/non-

carbonate components of both the dry-bulk and <20-mmfraction. The high concentration (wt%) of CaCO3 evaluatedon dry-bulk sediment (generally >60%) documents a minorinfluence from dilution by non-carbonate and terrigenousmaterial, especially in the lower portion of the studiedinterval. Therefore, CaCO3 fluctuations should reflect mainlysurface water carbonate production and export. Notably,carbonate content in the <20-mm fraction, representing thecalcareous nannofossil contribution to carbonate, corre-sponds closely to changes in bulk carbonate at this site,showing that CaCO3 is predominantly produced by cocco-liths. Beginning with MIS 6, however, increased discrepancybetween the two records testifies to a more consistentcontribution by other calcifying organisms, probablyforaminifera.Bulk CaCO3 flux (CARbulk; Fig. 5) is driven by the <20-mm

fraction (CAR<20mm; Fig. 5), confirming the role of coccolitho-phores as major carbonate producers. The contribution ofcoccolith-derived CaCO3, evaluated from quantitative count-ing (total CARLM; Fig. 5), fluctuates considerably from 0.06 to1.3 g cm�2 ka�1, with an average value of 0.33 g cm�2 ka�1

through the study interval. The two different methods appliedto evaluate the coccolith carbonate flux show similar trends,even if total CARLM documents fluxes that account for half ofthe CAR<20mm values (Fig. 5). However, note that combineduncertainties in ks and size are likely to yield a maximumerror of about 50% when species accumulation rates areconverted into specific flux rates (Ziveri et al., 2007). Addi-tional errors could also be introduced by counting problems,particularly if pellets or clumps of coccoliths are present. So,

Figure 3. Nannofossil dissolution index (NDI, from Bordiga et al.,2013), total coccolith concentration (Total Cc) and calcareousnannofossil distribution of principal taxa documented at ODP Site1209B. Total Cc is expressed in N g�1, while relative abundance oftaxa is expressed as a percentage. Grey bars show glacial intervals.

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COCCOLITHOPHORE CARBONATE IN THE NW PACIFIC OCEAN 61

considerable care should be taken in interpreting such data(Young and Ziveri, 2000). Moreover, several authors (e.g.Knappertsbusch et al., 1997; Sprengel et al., 2000;Baumann, 2004) have underlined that errors in coccolith-

carbonate calculations are substantial, particularly as a singlemean carbonate value does not account for the size variationsof most common species.Regarding total CARLM, along the MBE (MIS 12–9 interval

in our record) it is not higher than in the upper part of thecore, despite the high abundance of G. caribbeanica (Fig. 5).Worldwide, during this time interval, G. caribbeanica repre-sents the major dominant contributor to coccolithophoreassemblages (e.g. Bollmann et al., 1998; Barker et al., 2006),recording a very clear acme. For example, Baumann andFreitag (2004) found that CaCO3 accumulation more thandoubled during the acme in the Atlantic (Cape Basin). AtODP Site 1209B, this species characterized well the associa-tions but it is almost exclusively represented by specimens ofsmall size. This may explain why two relatively weak peaksof the total CARLM values (during MIS 12 and 9) are recordedacross the MBE interval. The total CARLM values coincidewith peaks in heavily calcified species, such as C. leptoporus,C. pelagicus, H. carteri, G. caribbeanica and G. muellerae.Conversely, very low values of total CARLM are recordedduring MIS 11.Maximum values of total CARLM are found at the MIS 8/

MIS 7 transition; decreased abundance of G. caribbeanica isfollowed by an abrupt increase of heavily calcified species.Here dominant calcite contributors are C. braarudii, C.leptoporus intermediate, C. pelagicus azorinus, C. pelagicuspelagicus, H. carteri, C. leptoporus small, C. quadriperforatusand G. oceanica medium. Another peak of total CARLM isobserved at the MIS 6/MIS 5 transition, caused by a newincrease of the same species cited above. Finally, the totalCARLM peak at the MIS 2/MIS 1 transition is related mainly toincreased carbonate production by a different morphotype ofC. leptoporus and by C. pelagicus azorinus, H. carteri and G.oceanica medium.At the glacial–interglacial scale, increases of total CARLM

occur during deglaciations at the MIS 8/MIS 7 and MIS 6/MIS5 transitions and to a lesser extent at the MIS 4/MIS 3transition; no particular long-term trend of total CARLM isevident during the last 450 ka (Fig. 5).

Time series analyses

The results of total NAR and total CARLM spectral-waveletanalyses (Fig. 6) show 58- and 90-ka periodicities. In the totalCARLM signal, the 58-ka cycle is stronger than the 90-kacycle; the latter periodicity is significant throughout therecord, whereas the 58-ka cycle signal is particularly strongfrom 380 ka upwards. In the total NAR signal, the 90-ka cycleis continuously recorded. The 90-ka cycle is recorded in bothsignals, and probably represents the response to eccentricityorbital forcing. Eccentricity is the main component of theEarth’s orbit affecting global annual insolation, which in turncontrols sea and atmospheric temperature, ultimately regulat-ing marine carbonate production.The 58-ka cycle, recorded in the total CARLM signal, could

be derived by the combined effects of obliquity and preces-sion orbital forcing variations (Berger, 1977). The effect ofthis orbital cycle combination is reflected in total insolationbalance controlling seasonal and latitudinal solar energyimpact and therefore local temperature changes.

Discussion

Calcareous nannoplankton carbonate dynamicsduring the last 450 ka

A recent study focusing on Holocene sediments suggestedthat roughly equal amounts of coccolith and foraminiferal

Figure 4. Primary productivity indices (N ratio and Sr/Ca ratio) andnannofossil accumulation rate (NAR). NAR is expressed inN cm�2 ka�1. Grey bars show glacial intervals.

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calcite rain to the open ocean seafloor (Broecker andClark, 2009). Here we investigate Pleistocene Pacific Oceansediments collected at depths <3000m to avoid dissolutionbias in coccolith versus foraminiferal calcite evaluation.PDW are very corrosive to calcium carbonate because theyrepresent an aged water mass with a light carbon isotopicsignature and high dissolved CO2 concentration. However, atpresent, ODP Site 1209B is bathed in waters saturated withrespect to calcite (V¼ 1.07 at 2322m; Berelson et al., 2007),and so is above the lysocline.Regarding coccolith production, before MIS 8 total NAR

(Fig. 7) displays maximum values related to the G. caribbean-ica acme during the MBE. Carbonate content (wt%) isparticularly high, and is dominated by coccolith calcite(Fig. 7): before MIS 8 coccoliths provide at least 60% of thecarbonate and as much as 90% during MIS 11 and 10. In thisinterval, the recorded highest values of primary productivity(N ratio, Fig. 7), documented also by Chiyonobu et al. (2012)for nearby Site 1210A, are consistent with modern regionalconditions during late winter/spring. During MIS 7, decreasedN ratio (Fig. 7) and increased F. profunda (Fig. 3) testify tolowering of primary productivity and strengthening of thethermocline, probably related to weakening of both thewesterly winds and the Kuroshio Extension, resemblingmodern summer conditions.Although total NAR decreases after MIS 7, Sr/Ca ratio

(Fig. 7) implies a proliferation of coccoliths, at least after thelatest part of MIS 7. From MIS 6 onward, total NAR values

remain low (see also the total coccolith concentration inFig. 3), and despite a consistently significant (>50% of bulkcarbonate) nannofossil contribution, CaCO3 (wt%) no longerreaches the values recorded before MIS 7. In contrast, the Nratio recovers. Short-term fluctuations in N ratio and totalNAR provide evidence of variability related to the glacial–interglacial cycles (Figs 6 and 7).Carbonate fluxes (total CARLM, CARbulk, CAR <20mm)

maintain constant mean values throughout the study interval(450 ka) with maxima coinciding with deglaciations andsubordinately with interglacial episodes (e.g. MIS 9) andminima during glacial phases. Total CARLM (Fig. 7), heredocumented for the NW Pacific, display comparable valuesto the south-west (Farrell and Prell, 1989) and eastern-equatorial Pacific (Lo�pez-Otalvaro et al., 2008). Conversely,higher values have been documented both for the North andSouth Atlantic (Oppo et al., 1990; Barker et al., 2006).Before MIS 8, the highest CaCO3 (wt%) values are almost

certainly related to coccolithophore production. From MIS 7onward, the observed changes in the coarse and fine CaCO3wt% (Fig. 7) probably represent changes in the relative contribu-tion of foraminifera and calcareous nannofossils to bulkcarbonate. Despite increased foraminifers, our data documentthat the coccolithophore calcite fraction largely dominatescarbonate sedimentation in the NW Pacific during the last 450ka, suggesting a high-productivity, coccolithophore-dominantand bloom-type ecological imprint; the foraminifera, instead,highlight a subordinate role in the carbonate export.

Figure 5. Distribution of carbonate and non-carbonate sediments (wt%) evaluated on both bulk and <20-mm fractions. Dry bulk density curve andcarbonate accumulation rate evaluated on bulk (CARbulk), on <20-mm fraction (CAR<20mm), and at the light microscope (Total CARLM). For lightmicroscope counting, CARLM from each nannofossil species is also reported and expressed in g cm�2 ka�1. Grey bars show glacial intervals.

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COCCOLITHOPHORE CARBONATE IN THE NW PACIFIC OCEAN 63

Moreover, from MIS 7 onward, decreased CaCO3 may alsobe due to increased dilution from terrigenous and otherbiogenic (diatoms) components (Fig. 7). In fact, increased inputof eolian Asian dust, diluting the carbonate component, andfavoring diatom productivity, has been documented by Zhanget al. (2007), particularly towards the end of MIS 7. Diatomscompete against coccoliths for nitrate and phosphate and arelimited by iron and silica supply (Goldman, 1984; Martinet al., 1994). During glacial times the eolian dust, rich in iron,increases (Fig. 7; Hovan et al., 1991; Zhang et al., 2007)leading to diatom proliferation. Today these biogenic fluxes ofcarbonate and opal follow seasonal patterns. High foraminiferfluxes are accompanied by high fluxes of organic matter andbiogenic opal during late spring and summer, with predictabledifferences in magnitude as a function of latitude (Mohiuddinet al., 2002, 2004; Eguchi et al., 2003).Decreased nannofossil abundance from MIS 7 onward at

Site 1209B, associated with increased biogenic opal andforaminifera observed at neighboring sites (Kawahata et al.,

1998; Yamane, 2003; Chiyonobu et al., 2012), underlines asignificant modification of the surface water mass towards themodern configuration after the MIS 7/8 transition (�243 ka inLR04), further amplifying a trend of increased North Pacificseasonality that began during the Plio-Pleistocene climatetransition (Haug et al., 2005; Venti and Billups, 2013). FromMIS 7 onward, intensification of the East Asian Monsoonprobably strengthened seasonality: warmer late-summer andautumn and colder winter–spring SSTs, causing alternationbetween intensely stratified and vertically mixed surfacewaters after 243 ka as also proposed by Chiyonobu et al.(2012) in this region of Site 1210A.Finally, wt% CaCO3 is also related to preservation on the

sea floor. Except for the earliest part of MIS 8 and the later partof MIS 7, NDI (Fig. 7, Bordiga et al., 2013) suggests mainlygood preservation and therefore over-saturation (V> 1) ofbottom water throughout the study interval. In fact, the MidBrunhes Dissolution interval (Hodell et al., 2000; Wanget al., 2003; Barker et al., 2006), recorded by the authors at

Figure 6. Power spectra and wavelet analyses of total CARLM and total NAR records processed through REDFIT and Lomb–Scargle analysis. Greenand black dotted lines in Lomb–Scargle power spectra indicate 95 and 80% confidence levels, respectively. Red line indicates the AR (1) theoreticalred noise spectrum (Schulz and Mudelsee, 2002). This figure is available in colour online at wileyonlinelibrary.com.

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MIS 11, is not recognizable at this site. Preservation highlightsa clear relationship with glacial–interglacial cycles, reflectingmainly, but not exclusively, the Pacific-style CaCO3 cycle (e.g.Naidu and Malmgren, 1999; Anderson et al., 2008; Sextonand Barker, 2012), where the calcium carbonate content ofdeep-sea sediments increases during glacial in comparisonwith interglacial intervals since 1.1 Ma (onset of the mid-Pleistocene Transition, see Clark et al., 2006).In this case, carbonate preservation on the sea floor is

strongly related to the aging effect of the bottom waters (Russonet al., 2009). The PDW being ‘more aged and corrosive’probably influenced our site during the interglacial phases. Weassume that during glacial intervals the boundary betweenPDW and the North Pacific Intermediate Water lowered due toincreased production of ‘cold and young’ intermediate waterfrom the Polar Front, so that our site was plunged in this latterwater, favoring carbonate preservation. A similar mechanismhas been invoked by Okazaki et al. (2012) for the last glacial inthe mid-latitude western North Pacific.In conclusion, changes in the ocean’s carbon and carbon-

ate export can affect ocean carbonate chemistry on time-scales between 1 and 100 ka, leading to shifts in the verticaldistribution of total dissolved inorganic carbon and totalalkalinity (Zeebe, 2012). Therefore, these processes areimportant to understand changes in ocean chemistry andatmospheric CO2 on glacial–interglacial timescales. Theabsence of any long-term trend in the CARLM record andresults of spectral and wavelet analyses on this time seriesillustrate that drivers >100 ka influence carbonate flux(Fig. 6). As changes in production and accumulation ofcarbonate in sedimentary reservoirs can drive significantchanges in pCO2 (e.g. Archer and Maier-Reimer, 1994;Sigman and Boyle, 2000), and coccolithophores potentiallycontribute 20–80% of this global carbonate fraction (e.g.Bornemann et al., 2003; Hay, 2004; Broecker and Clark,2009), the constant efficiency of the pelagic carbonate pumpcould explain the absence of any long-term trend in theatmospheric pCO2 recorded in the Antarctic ice-cores.

Paleoceanographic implications of dynamiccoccolith carbonate sedimentation during thelast 450 ka

The surface waters at our site are characterized by theSubtropical Front which separates the Subtropical Water mass(STW) to the south from the Transitional Water mass (TW) tothe north. At present, this oceanographic boundary migratesnorth and south seasonally. The TW is characterized by highmacronutrient concentration and the paleoproductivity andorganic matter fluxes decrease, from TW to STW, by nearlyan order of magnitude (Mohiuddin et al., 2002; Eguchi et al.,2003). Nannofossil assemblages and high values of primaryproductivity (except for MIS 7) document the influence of theTW over Site 1209 in almost the entire studied interval.During MIS 12–8, high productivity related to coccolitho-

phore blooming (see Sr/Ca ratio), good carbonate preserva-tion at the sea floor (seawater V calcite >1) as well as highSST (Fig. 7), documented by Ujiie (2003) and LaRiviere et al.(2012) for nearby sites (NP36; ODP 1208), are consistentwith the Kuroshio Extension in an ‘elongated state’ configura-tion, in which it has large meanders (Qiu, 2000). The pathand speed of the modern Kuroshio Current are interpreted tofluctuate in response to El Ni~no Southern Oscillation (ENSO)events (e.g. Teramoto and Kawabe, 1985). Modeling experi-ments suggest that ENSO-like paleoclimatic cyclicity mayhave also occurred in the past (Cane and Clement, 1999;Clement and Cane, 1999). Millennial-scale oscillations in

Figure 7. Summary of data collected at ODP Site 1209B (N, NDI,total NAR, CaCO3wt%, CARs, Sr/Ca), compared with eolian fluxes(Hovan et al., 1991), biogenic opal wt% (Kawahata et al., 1998;Yamane, 2003) and SST (Ujiie, 2003; LaRiviere et al., 2012) evaluatedin the area. SSTw, winter sea surface temperatures; SSTs, summer seasurface temperatures. Grey bars show glacial intervals.

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ENSO-like events have been documented in the Late Pleisto-cene successions (Pisias, 1978; Wang et al., 2000), indicatingthat they were pervasive in both time and space. Therefore,the Kuroshio Extension in an elongated state should causewarmer periods in the NW Pacific, offshore of Japan (Beaufortet al., 2001), as associated with La Ni~na events in theEquatorial Pacific region (Lo�pez-Otalvaro et al., 2008).During MIS 7, lowering of both paleoproductivity and SST

(Fig. 7) in the Shatsky Rise region seems consistent with theKuroshio Extension in a ‘contracted state’ configuration,characterized by weakening of meanders and persistence ofless productive subtropical waters over the site (e.g. Mohiud-din et al., 2004). Beaufort and Buchet (2003) reported a clearrelationship between the increase of F. profunda, recorded atour site during MIS 7, and an expanded warm pool in theWestern Pacific, as during modern El Ni~no conditions.During MIS 6 paleoproductivity increases again and the

coldest sea surface water conditions are recorded in theregion (Ujiie, 2003). Zhang et al. (2007) relate an episode ofhigh productivity during MIS 6 to enhanced input of dust fluxfrom Asia (Fig. 7). This mechanism for increased paleopro-ductivity is also consistent with ENSO perturbations andstrengthening of Boreal Monsoon dynamics (Zhang et al.,2007; Bordiga et al., 2013).The oceanographic conditions during MIS 5 mirror those

during MIS 11: therefore in the Pacific, hyperthermal episodesare associated with La Ni~na-like conditions.From the later part of the MIS 5 onward, paleoproductivity

seems relatively stable and SST documents a progressivedecrease (LaRiviere et al., 2012; Ujiie, 2003; Fig. 7), suggest-ing conditions unrelated to ENSO-like perturbations. Finally,the low resolution in the youngest part of the section preventsspeculation about the Holocene.

Concluding remarks

Coccolithophore records during the past 450 ka from theShatsky Rise highlight the dominant contribution of thiscalcifying functional group to carbonate sedimentation andits importance in carbon and carbonate export, as evidencedby N ratio and total NAR. Carbonate data on both the bulksediment and fine coccolith fraction indicate that nannofossilsare the main producers of pelagic carbonate, constituting 60–90% of the carbonate content of the sediments; foraminifera,by contrast, do not contribute significantly to carbonateproduction at the study site during the last 450 ka.Accumulation of coccolith-derived carbonates as deep-sea

sediments favors CO2 surplus buffering during carbonatedissolution episodes. The role of nannofossils in the carbonateexport and biological pump is particularly efficient from MIS12 to MIS 8 at the study site. Carbonate fluxes and dissolutionvary at glacial–interglacial scales; therefore, changes incoccolithophore production on timescales <100 ka influenceocean chemistry and atmospheric CO2 concentration.At the NW Pacific study site, coccolithophore assemblages

and primary productivity are driven by ocean–atmospherephenomena. La Ni~na-like conditions prevalent during theMBE favor coccolithophore proliferation and accumulation oftheir carbonate in sediments. By contrast, from MIS 6 onward,dust input from Asia during glacials favors diatom production.

Acknowledgements. We thank the Editor, Antony Long, and JorijntjeHenderiks for providing constructive comments that helped toimprove the manuscript. We are also grateful to two anonymousreviewers. We thank the Ocean Drilling Program for furnishing thesamples to M.C. that allowed the study of Site 1209B. This studyincorporates results of the PhD project of M.B. and was financially

supported by grants to M.C. and to P.Z. M. B. is grateful to the Instituteof Environmental Science and Technology of the Universitat Auton-oma de Barcelona and the Department of Analytical Chemistry atOviedo University, which made available laboratories and techniquesfor nannofossil flux evaluation and Sr/Ca analysis. Finally, weacknowledge financial support from Spanish Ministry of Science andInnovation (grant nos: CGL2009-10806; CTM2008-04365-E).

Abbreviations. CAR, coccolith accumulation rate; Cc, coccolithconcentration; ENSO, El Ni~no Southern Oscillation; MBE, Mid-Brunhes Event; MIS, marine isotope stage; NAR, nannofossil accumu-lation rate; NDI, nannofossil dissolution index; PDW, Pacific DeepWaters; SST, sea surface temperature; STW, Subtropical Water mass;TW, Transitional Water mass.

References

Aizawa C, Oba T, Okada H. 2004. Late Quaternary paleoceanog-raphy deduced from coccolith assemblages in a piston corerecovered off the central Japan coast. Marine Micropaleontology52: 277–297.

Anderson RF, Fleisher MQ, Lao Y, et al. 2008. Modern CaCO3

preservation in equatorial Pacific sediments in the context of late-Pleistocene glacial cycles. Marine Chemistry 111: 30–46.

Antonov JI, Seidov D, Boyer TP, et al. 2010. World Ocean Atlas 2009Vol. 2 Salinity. In NOAA Atlas NESDIS, Levitus S (ed). USGovernment Printing Office: Washington, DC.

Archer D, Maier-Reimer E. 1994. Effect of deep-sea sedimentarycalcite preservation on atmospheric CO2 concentration. Nature367: 260–263.

Barker S, Archer D, Booth L, et al. 2006. Globally increased pelagiccarbonate production during the mid-Brunhes dissolution intervaland the CO2 paradox of MIS 11. Quaternary Science Reviews 25:3278–3293.

Baumann K-H. 2004. Importance of size measurement forcoccolith carbonate flux estimates. Micropaleontology 50:35–43.

Baumann K-H, Bockel B, Frenz M. 2004. Coccolith contribution toSouth Atlantic carbonate sedimentation. In Coccolithophores fromMolecular Processes to Global Impact, Thierstein HR, Young J(eds). Springer: Berlin; 367–402.

Baumann K-H, Freitag T. 2004. Pleistocene fluctuations in thenorthern Benguela Current system as revealed by coccolithassemblages. Marine Micropaleontology 52: 195–215.

Beaufort L, Buchet N. 2003. Variability of specific coccolith CaCO3

weight, and primary productivity in the western pacific warm poolduring the last glacial cycle. Geophysical Research Abstracts(European Geosciences Union) 5: 13225.

Beaufort L, de Garidel-Thoron T, Mix AC, Pisias NG. 2001. ENSO-like forcing on oceanic primary production during the LatePleistocene. Science 293: 2440–2444.

Berelson WM, Balch WM, Najjar R, et al. 2007. Relating estimates ofCaCO3 production, export, and dissolution in the water column tomeasurements of CaCO3 rain into sediment traps and dissolutionon the sea floor: a revised global carbonate budget. GlobalBiogeochemical Cycles 21: DOI: 10.1029/2006GB002803.

Berger AL. 1977. Support for the astronomical theory of climaticchange. Nature 269: 44–45.

Berger WH, Winterer EL. 1974. Plate stratigraphy and the fluctuatingcarbonate line. In Pelagic Sediment: on Land and Under the Sea,International Sedimentologists Special Publication 1, Hsu KJ,Jenkyns HC (eds). Blackwell Publishing: Oxford; 11–48.

Bollmann J, Baumann K-H, Thierstein HR. 1998. Global dominanceof Gephyrocapsa coccoliths in the late Pleistocene: selectivedissolution, evolution, or global environmental change? Paleocea-nography 13: 517–529.

Bordiga M, Beaufort L, Cobianchi M, et al. 2013. Calcareousnannofossils as dissolution and primary productivity proxies in thelast 450 ka (Shatsky Rise, NW Pacific Ocean). Palaeogeography,Palaeoclimatology, Palaeoecology 371: 93–108.

Bornemann A, Aschwer U, Mutterlose J. 2003. The impact ofcalcareous nannofossils on the pelagic carbonate accumulation

Copyright # 2013 John Wiley & Sons, Ltd. J. Quaternary Sci., Vol. 29(1) 57–69 (2014)

66 JOURNAL OF QUATERNARY SCIENCE

across the Jurassic–Cretaceous boundary. Palaeogeography, Palae-oclimatology, Palaeoecology 199: 187–228.

Bown PR. 1998. Calcareous Nannofossil Biostratigraphy. Chapman &Hall: Cambridge.

Boyd PW, Watson AJ, Law CS, Abraham ER, Trull T, Murdoch R,Bakker DC, Bowie AR, Buesseler KO, Chang H, Charette M, CrootP, Downing K, Frew R, Gall M, Hadfield M, Hall J, Harvey M,Jameson G, LaRoche J, Liddicoat M, Ling R, Maldonado MT,McKay RM, Nodder S, Pickmere S, Pridmore R, Rintoul S, Safi K,Sutton P, Strzepek R, Tanneberger K, Turner S, Waite A, Zeldis J.2000. A mesoscale phytoplankton bloom in the polar SouthernOcean stimulated by iron fertilization. Nature 407: 695–702.

Bralower TJ, Premoli Silva I, Malone MJ, et al. 2002. Site 1209. InProceedings Ocean Drilling Program, Initial Reports, 198, BralowerTJ, Premoli Silva I, Malone MJ et al. (eds). ODP: College Station,TX; 1–102.

Broecker JS. 1995. Chaotic climate. Scientific American 273:62–68.

Broecker W, Clark E. 2009. Ratio of coccolith CaCO3 to foraminiferaCaCO3 in late Holocene deep sea sediments. Paleoceanography24: DOI: 10.1029/2009PA001731.

Cane MA, Clement A. 1999. A role for the tropical Pacific coupledocean-atmosphere system on Milankovitch and millennial timescales. Part II: Global impacts. In Mechanisms of Global ClimateChange at Millennial Time-Scales, Geophysical Monograph Series,Clark PU, Webb RS, Keigwin LD (eds). American GeophysicalUnion: Washington, DC; 373–383.

Chai F, Jiang M, Barber RT, et al. 2003. Interdecadal variation of theTransition Zone chlorophyll front, a physical-biological modelsimulation between 1960 and 1990. Journal of Oceanography 59:461–475.

Chiu TC, Broecker WS. 2008. Toward better paleocarbonate ionreconstructions: New insights regarding the CaCO3 size index.Paleoceanography 23: PA2216. doi: 10.1029/2008PA001599.

Chiyonobu S, Mori Y, Oda M. 2012. Reconstruction of paleoceano-graphic conditions in the northwestern Pacific Ocean over the last500 kyr based on calcareous nannofossil and planktic foraminiferalassemblages. Marine Micropaleontology 96-97: 29–37.

Clark PU, Archer D, Pollard D. 2006. The middle Pleistocenetransition: characteristics, mechanisms, and implications for long-term changes in atmospheric pCO2. Quaternary Science Reviews25: 3150–3184.

Clement A, Cane MA. 1999. A role for the tropical Pacific coupledocean-atmosphere system on Milankovitch and millennial timescales. Part I: A modeling study of tropical Pacific variability. InMechanisms of Global Climate Change at Millennial Time-Scales,Geophysical Monograph Series, Clark PU, Webb RS, KeigwinLD (eds). American Geophysical Union: Washington, DC; 363–371.

Da Silva A, Young-Molling AC, Levitus S. 1994. Atlas of SurfaceMarine Data 1994, vol. 3: Anomalies of Fluxes of Heat andMomentum. NOAA Atlas NESDIS 8. U.S. Gov. Printing Office:Washington, DC; 413 pp.

Eguchi NO, Ujiie H, Kawahata H, et al. 2003. Seasonal variations inplanktonic foraminifera at three sediment traps in the Subarctic,Transition and Subtropical zones of the central North PacificOcean. Marine Micropaleontology 48: 149–163.

Farrell JW, Prell WL. 1989. Climatic change and CaCO3 preservationan 800,000 year bathymetric reconstruction from the centralequatorial Pacific Ocean. Paleoceanography 4: 447–466.

Flores JA, Gersonde R, Sierro FJ, et al. 2000. Southern Pleistocenecalcareous nannofossil events. Calibrations with isotope andgeomagnetic stratigraphies. Marine Micropaleontology 40: 377–402.

Foster G. 1996a. Time series analysis by projection. I. Statisticalproperties of Fourier analysis. Astronomical Journal 111: 541–554.

Foster G. 1996b. Time series analysis by projection. II. Tensormethods for time series analysis. Astronomical Journal 111: 555–566.

Foster G. 1996c. Wavelets for period analysis of unevenly sampledtime series. Astronomical Journal 112: 1709–1729.

Geisen M, Billard C, Broese ATC, et al. 2002. Life-cycle associationsinvolving pairs of holococcolithophorid species: intraspecific varia-

tion or cryptic speciation? European Journal of Phycology 37: 531–550.

Goldman JC. 1984. Oceanic nutrient cycles. In Flows of Energy andMaterials in Marine Ecosystems: Theory and Practice, NATOConference Series Volume 13, Fasham MJR (ed). Springer-Verlag:New York; 137–170.

Harada N, Sato M, Shiraishi A, et al. 2006. Characteristics ofalkenone distributions in suspended and sinking particles in thenorthwestern North Pacific. Geochimica et Cosmochimica Acta70: 2045–2062.

Haug GH, Ganopolski A, Sigman DM, Rosell-Mele A, Swann GE,Tiedemann R, Jaccard SL, Bollmann J, Maslin MA, Leng MJ,Eglinton G. 2005. North Pacific seasonality and the glaciation ofNorth America 2.7 million years ago. Nature 433: 821–825.

Haug GH, Maslin MA, Sarnthein M, et al. 1995. Evolution ofnorthwest Pacific sedimentation patterns since 6 Ma (Site 882). InProceedings of Ocean Drilling Program, Scientific Results, 145,Rea DK, Basov IA, Scoll DW, Allan F (eds). ODP: College Station,TX; 293–301.

Hay WW. 2004. Carbonate fluxes and calcareous nannoplankton. InCoccolithophores: from Molecular Processes to Global Impact,Thierstein H, Young J (eds). Springer Verlag: Berlin; 509–528.

Hine N, Weaver PPE. 1998. Quaternary. In Calcareous NannofossilBiostratigraphy, Bown PR (ed). Chapman & Hall: Cambridge; 266–283.

Hodell DA, Charles CD, Ninnemann US. 2000. Comparison ofinterglacial stages in the South Atlantic sector of the southernocean for the past 450 kyr: implifications for Marine Isotope Stage(MIS)11. Global and Planetary Change 24: 7–26.

Hovan SA, Rea DK, Pisias NG. 1991. Late Pleistocene continentalclimate and oceanic variability recorded in northwest Pacificsediments. Paleoceanography 6: 349–370.

Incarbona A, Ziveri P, Sabatino N, et al. 2011. Conflicting coccoli-thophore and geochemical evidence for productivity levels in theEastern Mediterranean Sapropel S1. Marine Micropaleontology 81:131–143.

Jansen JHF, Kuijpers A, Troelstra SR. 1986. A mid-Brunhes climaticevent: long-term changes in global atmosphere and ocean circula-tion. Science 232: 619–622.

Jordan RW, Cros L, Young JR. 2004. A revised classification schemefor living haptophytes. Micropaleontology 50: 55–79.

Kalnay E, Kanamitsu M, Kistler R, et al. 1996. The NCEP/NCAR 40-Year Reanalysis Project. Bulletin of the American MeteorologySociety 77: 437–471.

Kawahata H, Ohshima H. 2002. Small latitudinal shift in theKuroshio Extension (Central Pacific) during glacial times: evidencefrom pollen transport. Quaternary Science Reviews 21: 1705–1717.

Kawahata H, Okamoto T, Matsumoto E, et al. 2000. Fluctuations ofeolian flux and ocean productivity in the midlatitude north Pacificduring the last 200 kyr. Quaternary Science Reviews 19: 1279–1291.

Kawahata H, Suzuki A, Ahagon N. 1998. Biogenic sediments in theWest Caroline Basin, the equatorial Pacific during the last 330,000years. Marine Geology 149: 155–176.

Keigwin LD. 1987. North Pacific deep-water formation during thelatest glaciation. Nature 330: 362–364.

Klaas C, Archer D. 2003. Spatial patterns of organic carbon, opaland CaCO3 fluxes in sediment traps: application to the LGMcarbon cycle. AGU Fall Meeting Abstract. Nice, France; abstract#9982.

Knappertsbusch M, Cortes MY, Thierstein HR. 1997. Morphologicvariability of the coccolithophorid Calcidiscus leptoporus(Murray and Blackman) Loeblich and Tappan during the Holoceneand early Pleistocene. Marine Micropaleontology 30: 293–317.

Kuroyanagi A, Kawahata H, Nishi H, et al. 2002. Seasonal changes inplanktonic foraminifera in the northwestern North Pacific Ocean:sediment trap experiments from subarctic and subtropical gyres.Deep-Sea Research Part II 49: 5627–5645.

LaRiviere JP, Ravelo AC, Crimmins A, Dekens PS, Ford HL, Lyle M,Wara MW. 2012. Late Miocene decoupling of oceanic warmth andatmospheric carbon dioxide forcing. Nature 486: 97–100.

Copyright # 2013 John Wiley & Sons, Ltd. J. Quaternary Sci., Vol. 29(1) 57–69 (2014)

COCCOLITHOPHORE CARBONATE IN THE NW PACIFIC OCEAN 67

Levitus S, Boyer T. 1994. World Ocean Atlas 1994, Vol. 4:Temperature. NOAA Atlas NESDIS 4, U.S. Gov. Printing Office:Washington, DC; 117 pp.

Lisiecki LE, Raymo ME. 2005. A Plio-Pleistocene stack of 57 globallydistributed benthic d18O records. Paleoceanography 20: DOI:10.1029/2004PA001071.

Lo�pez-Otalvaro G-E, Flores J-A, Sierro FJ, et al. 2008. Variations incoccolithophorid production in the Eastern Equatorial Pacific atODP Site 1240 over the last seven glacial–interglacial cycles.Marine Micropaleontology 69: 52–69.

Lototskaya A, Ziveri P, Ganssen GM, et al. 1998. Calcareousnannoflora response to Termination II at 45˚ N, 25˚ W (northeastAtlantic). Marine Micropaleontology 34: 47–70.

Marino M, Maiorano P, Lirer F, et al. 2009. Response of calcareousnannofossil assemblages to paleoenvironmental changes throughthe mid-Pleistocene revolution at Site 1090 (Southern Ocean).Palaeogeography, Palaeoclimatology, Palaeoecology 280: 333–349.

Martin JH. 1990. Glacial–interglacial CO2 change: the iron hypothe-sis. Paleoceanography 5: 1–13.

Martin JH, Coale KH, Johnson KS, et al. 1994. Testing the ironhypothesis in ecosystems of the equatorial Pacific Ocean. Nature371: 123–129.

Millero FJ. 1996. Chemical Oceanography. CRC Press: Boca Raton.Mizuno K, White WB. 1983. Annual and interannual variability inthe Kuroshio Current system. Journal of Physical Oceanography 13:1847–1867.

Mohiuddin MM, Nishimura A, Tanaka Y, et al. 2002. Regional andinterannual productivity of biogenic components and planktonicforaminiferal fluxes in the northwestern Pacific Basin. MarineMicropaleontology 45: 57–82.

Mohiuddin MM, Nishimura A, Tanaka Y, et al. 2004. Seasonality ofbiogenic particle and planktonic foraminifera fluxes: response tohydrographic variability in the Kuroshio Extension, northwesternPacific Ocean. Deep-Sea Research Part I 51: 1659–1683.

Muller G, Gastner M. 1971. The ‘Karbonate-Bombe’, a simple devicefor the determination of the carbonate content in sediments, soilsand other materials. Neues Jahrbuch fur Mineralogie, Monatshefte10: 466–469.

Naidu PD, Malmgren BA. 1999. Quaternary carbonate record fromthe equatorial Indian Ocean and its relationship with productivitychanges. Marine Geology 161: 49–62.

Noriki S, Tsunogai S. 1986. Particulate fluxes and major componentsof settling particles from sediment trap experiments in the PacificOcean. Deep Sea Research. Part A, Oceanographic ResearchPapers 33: 903–912.

Okazaki Y, Sagawa T, Asahi H, et al. 2012. Ventilation changes inthe western North Pacific since the last glacial period. Climate ofthe Past 8: 17–24.

Oppo D, Fairbanks R, Gordon AL, et al. 1990. Late PleistoceneSouthern Ocean d13C variability. Paleoceanography 5: 43–54.

Palike H, Mitchell WL, Nishi H, et al. 2012. A Cenozoic record ofthe equatorial Pacific carbonate compensation depth. Nature 488:609–614.

Parente A, Cachao M, Baumann K-H, et al. 2004. Morphometry ofCoccolithus pelagicus s.l. (Coccolithophore, Haptophyta) fromoffshore Portugal, during the last 200 kyr. Micropaleontology 50:107–120.

Pisias NG. 1978. Paleoceanography of the Santa Barbara Basinduring the last 8000 years. Quaternary Research 10: 366–384.

Qiu B. 2000. Interannual variability of the Kuroshio Extension systemand its impact on the wintertime SST field. Journal of PhysicalOceanography 30: 1486–1502.

Qiu B. 2001. Kuroshio and Oyashio currents. In Encyclopedia ofOcean Science. Second Edition, Steele JH, Turekian KK, et al.(eds). Reference Module in Earth Systems and EnvironmentalSciences, Elsevier: 358–369.

Rickaby REM, Bard E, Sonzogni C, et al. 2007. Coccolith chemistryreveals secular variations in the global ocean carbon cycle? Earthand Planetary Sciences Letters 253: 83–95.

Ridgwell A, Zeebe RE. 2005. The role of the global carbonate cyclein the regulation and evolution of the Earth system. Earth andPlanetary Sciences Letters 234: 299–315.

Russon T, Elliot M, Kissel C, et al. 2009. Middle-late Pleistocene deepwater circulation in the southwest subtropical Pacific. Paleocea-nography 24: DOI: 10.1029/2009PA001755.

Sancetta C, Silvestri SM. 1986. Pliocene–Pleistocene evolution of thenorth Pacific ocean–atmosphere system, interpreted from fossildiatoms. Paleoceanography 1: 163–180.

Schulz M, Mudelsee M. 2002. REDFIT: estimating red-noise spectradirectly from unevenly spaced paleoclimatic time series. Com-puters and Geosciences 28: 421–426.

Sexton PF, Barker S. 2012. Onset of ‘Pacific-style’ deep-sea sedimen-tary carbonate cycles at the mid-Pleistocene transition. Earth andPlanetary Sciences Letters 321-322: 82–94.

Sigman DM, Boyle EA. 2000. Glacial/interglacial variations inatmospheric carbon dioxide. Nature 407: 859–869.

Sprengel C, Baumann K-H, Nauer S. 2000. Seasonal and interannualvariation of coccolithophore fluxes and species composition insediment traps north of Gran Canaria (29˚ N 15˚ W). MarineMicropaleontology 39: 157–178.

Stoll HM, Schrag DP. 2000. Coccolith Sr/Ca as a new indicator ofcoccolithophorid calcification and growth rate. Geochemistry,Geophysics, Geosystems 1: 1–24.

Stoll HM, Schrag DP. 2001. Sr/Ca variations in Cretaceous carbonate:relation to productivity and sea level changes. Palaeogeography,Palaeoclimatology, Palaeoecology 168: 311–336.

Stoll HM, Ziveri P, Geisen M, et al. 2002. Potential and limitations ofSr/Ca ratios in coccolith carbonate: new perspectives from culturesand monospecific samples from sediments. Philosopical Trans-actions of the Royal Society, London 360: 719–747.

Teramoto T, Kawabe M. 1985. Current and topographical effect inthe Kuroshio Current zone. In Oceanic Characteristics and TheirChanges, Kajiura K (ed). Kouseisha-Kouseikaku: Tokyo; 79–88 (inJapanese).

Sarmiento JL, Le Quere C. 1996. Oceanic carbon dioxide uptake in amodel of century-scale global warming. Science 271: 1346–1350.

Thierstein HR, Geitzenauer KR, Molfino B. 1977. Global synchrone-ity of the late Quaternary coccolith datum levels: validation byoxygen isotopes. Geology 5: 400–404.

Ujiie H. 2003. A 370-ka paleoceanographic record from the HessRise, central North Pacific Ocean, and an indistinct ‘KuroshioExtension’. Marine Micropaleontology 49: 21–47.

Venti NL, Billups K. 2013. Surface water hydrography of the KuroshioExtension during the Pliocene–Pleistocene climate transition. Ma-rine Micropaleontology 101: 106–114.

Vincent E. 1975. Neogene planktonic foraminifera from thecentral north pacific, leg 32, Deep Sea Drilling Project. In InitialReports of the Deep Sea Drilling Project 32, Larson RL, Moberly R,et al. (eds). Government Printing Office: Washington, DC; 765–801.

Wang H, Follmer LR, Liu JC. 2000. Isotope evidence of paleo-ElNi~no-Southern Oscillation cycles in loess-paleosol record in thecentral United States. Geology 28: 771–774.

Wang P, Tian J, Cheng X, et al. 2003. Carbon reservoir changespreceded major ice-sheet expansion at the mid-Brunhes event.Geology 31: 239–242.

Warren BA. 1983. Why is no deep water formed in the North Pacific.Journal of Marine Research 41: 327–347.

Westbroek P, Brown CW, van Bleijswijk J, et al. 1993. A modelsystem approach to biological climate forcing. The example ofEmiliania huxleyi. Global and Planetary Change 8: 27–46.

Wilson C, Villareal TA, Brzezinski MA, Krause JW, Shcherbina AY.2013. Chlorophyll bloom development and the subtropical front inthe North Pacific. Journal of Geophysical Research, Oceans 118:1473–1488.

Xie P, Arkin PA. 1997. Global precipitation: 17-year monthly analysisbased on gague observations, satellite estimates, and numericaloutputs. Bulletin of the American Meteorology Society 78: 2539–2558.

Yamane M. 2003. Late Quaternary variations in water mass in theShatsky Rise area, northwest Pacific Ocean. Marine Micropaleon-tology 48: 205–223.

Young JR. 1998. Neogene. In Calcareous Nannofossil Biostratigraphy,Bown PR, (ed). Chapman & Hall: Cambridge; 225–265.

Copyright # 2013 John Wiley & Sons, Ltd. J. Quaternary Sci., Vol. 29(1) 57–69 (2014)

68 JOURNAL OF QUATERNARY SCIENCE

Young JR, Bown PR, Lees JA. 2011. Nannotax Website. InternationalNannoplankton Association, http://nannotax.org. [accessed 1 January2010].

Young JR, Geisen M, Cros L, et al. 2003. A guide to extantcoccolithophore taxonomy. Journal of Nannoplankton Research,Special Issue 1.

Young JR, Ziveri P. 2000. Calculation of coccolith volume and its usein calibration of carbonate flux estimates. Deep-Sea Research PartII 47: 1679–1700.

Zeebe RE. 2012. History of seawater carbonate chemistry, atmosphericCO2 and ocean acidification. Annual Review of Earth PlanetarySciences 40: 141–165.

Zhang J, Wang P, Li Q, et al. 2007. Western equatorial Pacificproductivity and carbonate dissolution over the last 550 kyr:foraminiferal and nannofossil evidence from ODP Hole 807A.Marine Micropaleontology 64: 121–140.

Ziveri P, de Bernardi B, Baumann K-H, et al. 2007. Sinking ofcoccolith carbonate and potential contribution to organic carbonballasting in the deep ocean. Deep-Sea Research Part II 54: 659–675.

Ziveri P, Young JR, van Hinte JE. 1999. Coccolithophore exportproduction and accumulation rates. GeoResearch Forum 5:41–56.

COCCOLITHOPHORE CARBONATE IN THE NW PACIFIC OCEAN 69

Copyright # 2013 John Wiley & Sons, Ltd. J. Quaternary Sci., Vol. 29(1) 57–69 (2014)