Reprint - Gvsu - Grand Valley State University

Reprint - Gvsu - Grand Valley State University

ClickHereforFullArticlePALEOCEANOGRAPHY, VOL. 23, PA1216, doi:10.1029/2007PA001433, 2008Effects of surface ocean conditions on deep-sea calcite dissolutionproxies in the tropical PacificFigen Mekik 1 and Lisa Raterink 2Received 14 February 2007; revised 23 October 2007; accepted 8 November 2007; published 26 March 2008.[1] Finding the ideal deep-sea CaCO 3 dissolution proxy is essential for quantifying the role of the marinecarbonate system in regulating atmospheric pCO 2 over millennia. We explore the potential of using theGloborotalia menardii fragmentation index (MFI) and size-normalized foraminifer shell weight (SNSW) ascomplementary indicators of deep-sea CaCO 3 dissolution. MFI has strong correlations with bottom water[CO 2 3 ], modeled estimates of percent CaCO 3 dissolved, and Mg/Ca in Pulleniatina obliquiloculata in core topsamples along a depth transect on the Ontong Java Plateau (OJP) where surface ocean temperature variation isminimal. SNSW of P. obliquiloculata and Neogloboquadrina dutertrei have weak correlations with MFI-basedpercent dissolved, Mg/Ca in P. obliquiloculata shells and bottom water [CO 23 ] on the OJP. In core top samplesfrom the eastern equatorial Pacific (EEP), SNSW of P. obliquiloculata has moderate to strong correlationswith both MFI-based percent CaCO 3 dissolved estimates and surface ocean environmental parameters. SNSWof N. dutertrei shells shows a latitudinal distribution in the EEP and a moderately strong correlation withMFI-based percent dissolved estimates when samples from the equatorial part of the region are excluded. Ourresults suggest that there may potentially be multiple genotypes of N. dutertrei in the EEP which may bereflected in their shell weight. MFI-based percent CaCO 3 dissolved estimates have no quantifiable relationshipwith any surface ocean environmental parameter in the EEP. Thus MFI acts as a reliable quantitative CaCO 3dissolution proxy insensitive to environmental biases within calcification waters of foraminifers.Citation: Mekik, F., and L. Raterink (2008), Effects of surface ocean conditions on deep-sea calcite dissolution proxies in the tropicalPacific, Paleoceanography, 23, PA1216, doi:10.1029/2007PA001433.1. Introduction[2] Accurately quantifying deep marine CaCO 3 dissolutionhas been a challenging oceanographic problem formany decades [e.g., Arrhenius, 1952; Berger, 1973;Broecker, 1982; Archer and Maier-Reimer, 1994; Mekikand François, 2006]. Calcite preservation is a major componentof the marine carbonate system (others include theinflux of ions into the ocean as weathering products fromland, the air-sea exchange of CO 2 , the marine biologicalpump , and the rain ratio, which is the ratio of organic carbonto calcite flux at the seabed); and developing a reliableCaCO 3 preservation proxy is important because the dissolutionof carbonates in deep-sea sediments is an integralpart of the global carbon cycle in regulating atmosphericpCO 2 over thousands of years [Broecker, 1971; Archer andMaier-Reimer, 1994; Archer et al., 2000].[3] Most CaCO 3 dissolution indicators are based, at leastin part, on the preservation state of foraminifer shells. Someof the ways this preservation state has been defined are (1)the ratio of the number of foraminifer test fragments for agiven species to the number of whole shells from thatspecies [e.g., Peterson and Prell, 1985a, 1985b; Le and1 Department of Geology, Grand Valley State University, Allendale,Michigan, USA.2 Department of Earth and Environmental Sciences, Wright StateUniversity, Dayton, Ohio, USA.Copyright 2008 by the American Geophysical Union.0883-8305/08/2007PA001433$12.00Shackleton, 1992; Mekik et al., 2002]; (2) dissolutioninducedloss in size-normalized whole foraminifer shellweight [Lohmann, 1995; Broecker and Clark, 2001a,2001b]; and (3) changes in the Mg/Ca ratio of foraminifershells through dissolution [Brown and Elderfield, 1996;Rosenthal et al., 2000; Dekens et al., 2002; Rosenthal andLohmann, 2002; Mekik and François, 2006; Mekik et al.,2007a]. Most proxies anchor dissolution-induced changes inforaminifer shells to the [CO 23 ] of bottom waters [e.g.,Broecker and Clark, 2001a, 2001b; Dekens et al., 2002;Marchitto et al., 2005]. Instead, Mekik et al. [2002] relatedthe fragmentation trend of Globorotalia menardii shells tomodel derived estimates of percent CaCO 3 dissolved indeep-sea sediments. Their percent CaCO 3 dissolved esti-2mates take into account both the CO 3 undersaturation ofbottom waters and respiratory CaCO 3 dissolution withinsediments driven by fluxes of organic carbon reaching theseabed [Emerson and Bender, 1981].[4] The ideal CaCO 3 dissolution proxy would be (1) timeefficient(short analysis time per sample); (2) based onspecies (and/or their fragments) which are easy to identifyeven by nonspecialists, if dependent on biogenic components;(3) without biological/ecological bias, or at leasthave a bias that is quantifiable and accurately predictable;(4) sensitive to a wide range of dissolution (ideally from 0 to100% calcite dissolved); (5) calibrated against an independentand quantitative estimate of percent CaCO 3 dissolved;and (6) reliably applicable in areas with strong gradients tosurface ocean conditions like temperature, [CO 3 2 ], nutrientPA12161of15

PA1216MEKIK AND RATERINK: CALCITE DISSOLUTION IN THE DEEP SEAPA1216availability, productivity, and even in areas where there islarge variation to both organic carbon and calcite fluxesreaching the seabed.[5] We assume uniformity among foraminifer shells whendeveloping CaCO 3 dissolution proxies, yet no two foraminifersare exactly alike because vital processes may affectshell composition, thickness and therefore weight. The bestmeans for quantifying deep-sea CaCO 3 dissolution may beusing a multiproxy approach within the same sedimentsamples [e.g., Mekik and François, 2006; Naik and Naidu,2007; Ni et al., 2007]. We will focus on two dissolutionproxies, size normalized whole foraminifer shell weight(SNSW) [Broecker and Clark, 2001a; 2001b, 2003] andthe G. menardii fragmentation index (MFI) [Mekik et al.,2002], with some independent corroboration from foraminiferMg/Ca. We undertake the following research questions:[6] 1. Intuitively, loss in foraminifer shell weight wouldprecede fragmentation as dissolution progresses. Is theresuch a sequential relationship between SNSW and MFI withincreasing CaCO 3 dissolution in the sediments and decreasingbottom water [CO 23 ]? Or does shell weight loss happensimultaneously with fragmentation under similar degrees of2bottom water CO 3 undersaturation and organic carbondegradation in sediment pore waters?[7] 2. How sensitive are SNSW and MFI to environmentalinfluences within foraminifers’ calcification waters, such astemperature, [CO 23 ], nutrient availability and apparent oxygenutilization (AOU)? Mg/Ca in foraminifer shells is predominantlygoverned by calcification temperature [e.g.,Nürnberg, 1995; Elderfield and Ganssen, 2000; Lea et al.,2000; Anand et al., 2003]; is this also true of SNSWand MFI?2. Background2.1. Globorotalia menardii Fragmentation Index[8] The G. menardii fragmentation index is the ratio ofthe number of damaged G. menardii specimens (D) to thenumber of whole plus damaged specimens of this specieswithin a sediment aliquot. Damaged specimens are groupedinto categories as whole specimens with small holes (holes),pieces greater than half intact (>half), pieces less than halfintact ( halfþ ðnumber < half=3Þ þ ðnumber keels=5Þ ð1Þ[9] Mekik et al. [2002] based MFI on Ku and Oba’s[1978] laboratory experiments, which showed that dissolutiondamage in G. menardii shells is quantifiable. Theavailable MFI transfer function relates the fragmentationtrend of G. menardii shells in core tops of deep Pacificsediments to model-derived estimates of percent CaCO 3dissolved (R 2 = 0.88) with the following calibration equation[Mekik et al., 2002]:percent CaCO 3 dissolved ¼5:111 þ ðMFI*160:491ÞMFI 2 *79:636 ð2Þ[10] By percent CaCO 3 dissolved, we mean the fraction ofthe vertical calcite flux that has been lost to dissolution inany one spot on the sea bottom. Mekik et al. [2002] used thebiogeochemical model Muds [Archer et al., 2002] to calculatethe percent CaCO 3 dissolved for sample locations alongtwo depth transects in the Pacific Ocean: on the OntongJava Plateau (OJP), and on the East Pacific Rise outside ofthe equatorial upwelling region (1900–4441 m depth).These values were then used to calibrate MFI. Both bottom2water DCO 3 (which is the [CO 23 ] of in situ waters less[CO 23 ] at saturation) and organic carbon fluxes reachingthe sediments were included both in the model [Archer etal., 2002] and in calculations of percent CaCO 3 dissolved[Mekik et al., 2002]. This is because CaCO 3 dissolution onthe seafloor is in part driven by organic carbon degradationin the top meter of sediment.[11] All calibration samples experienced some bottom22water CO 3 undersaturation (DCO 3 ranges between20.32 and 28.98 mmol/kg). Estimates of DCO 3 for eachsample location in MFI’s calibration sample set are fromArcher’s [1996, personal communication, 2001] globalgridded database. Organic carbon flux estimates used tocalibrate MFI are from (1) satellite-based surface oceanproductivity estimates from Behrenfeld and Falkowski[1997]; (2) surface ocean productivity compilations ofBerger et al. [1987] and Berger [1989] and the attenuationof organic carbon with water depth using Berger et al.’s[1987] equation; and (3) Jahnke’s [1996] global griddeddatabase for benthic oxygen fluxes. Details regarding theequation, calibration and modeling of MFI are discussed byMekik et al. [2002].[12] Mekik and François [2006] provided independentcorroboration for MFI as a dissolution proxy using Mg/Caand Mg/Sr in shells of P. obliquiloculata and G. menardii insamples from the OJP where surface ocean temperaturevariation is minimal but where there is a steep gradient to2DCO 3 of bottom waters. Mg/Ca in P. obliquiloculatashells strongly correlates with MFI (R 2 = 0.94) and withMFI-based percent dissolved (R 2 = 0.84) on the OJP [Mekikand François, 2006] where Mg/Ca decreases with increasingMFI-based percent calcite dissolved. Subsequently,Mekik et al. [2007a] used MFI for dissolution correctionof Mg/Ca paleothermometry. Mekik et al. [2002, 2007b]expanded MFI’s applicability to core top samples in theeastern equatorial Pacific (EEP) where both surface oceanproductivity and the rain ratio reaching the seabed arehighly variable. Also, Loubere et al. [2004] and Richaudet al. [2007] applied MFI in down core work for estimatingCaCO 3 fluxes to the deep sea.[13] In summary, MFI is unique among available dissolutionproxies because (1) G. menardiis provide a quantifiablefragmentation trend with increasing dissolution whereother species tend to stay intact until a threshold value of2DCO 3 is reached, and then fall to pieces randomly belowthis threshold (F. Mekik, unpublished data, 2000); (2) it isthe only dissolution proxy anchored against model-derivedestimates of percent CaCO 3 dissolved per sample location[Mekik et al., 2002]; (3) it is efficient (20–30 minutes persample); (4) it uses a species whose fragments are easy toidentify; (5) it works at least in one region (EEP) where the2of15

PA1216MEKIK AND RATERINK: CALCITE DISSOLUTION IN THE DEEP SEAPA1216surface ocean has a strong productivity gradient [Mekik etal., 2002, 2007a, 2007b]; and (6) as explained above, thereis some independent corroboration for MFI as a dissolutionproxy from Mg/Ca and Mg/Sr in multiple species ofplanktonic foraminifers [Mekik and François, 2006]. However,neither the relationship between MFI and SNSW, norSNSW’s application in an upwelling region like the EEPhas previously been explored. That is our goal herein.2.2. Size-Normalized Foraminifer Shell Weight[14] The size-normalized foraminifer shell weight(SNSW) method is founded on the assumption that foraminifertest weight loss within a specified size range isdriven solely by dissolution of foraminifer shells in sediments[Lohmann, 1995; Broecker and Clark, 2001a,2001b]. This has been well established for several speciesof planktonic foraminifers including Neogloboquadrinadutertrei, Pulleniatina obliquiloculata and Globigerinoidesruber [e.g., Broecker and Clark, 2001a, 2001b, 2003].Broecker and Clark [2001a] relate shell mass loss to depthnormalized bottom water [CO 23 ] which they define asCO 2 3 * ¼ CO 2 3 þ 20ð4 zÞð3Þwhere [CO 2 3 ]* represents depth normalized [CO 2 3 ] and zis water depth in kilometers. [CO 23 ] values in their workare extrapolated from GEOSECS data. They report anaverage size normalized foraminifer weight loss slope of0.30 ± 0.05 mg per 1 mmol/kg decrease in depth normalized[CO 2 3 ].[15] It seems that SNSW and MFI may potentially serveto expand and complement one another since, intuitively,foraminifer shell mass loss should precede fragmentation.We explore this issue as well as potential environmentaleffects on each proxy.2.3. Eastern Equatorial Pacific[16] Unlike surface waters above the Ontong Java Plateau,the EEP is an expansive region of both coastal andequatorial upwelling with high pCO 2 in its surface waters[Tans et al., 1990]. The South Equatorial Current (SEqC) isdriven by trade winds and marks the northern branch of theSouth Pacific subtropical gyre [Pennington et al., 2006]where it feeds a major open ocean upwelling system in theEEP. The SEqC seems to originate from the SW AntarcticPacific [Toggweiler et al., 1991; Kessler, 2006]. The EEPcold tongue results from the divergence of flow along theequator and generally spans between 3°N and 3°S though itis not usually symmetrical about the equator [Wyrtki, 1981;Fiedler and Talley, 2006]. This cold upwelling processbrings macronutrients to the euphotic zone [Chavez andBarber, 1987] and the deep chlorophyll maximum is shallowin this region of the EEP [Fiedler and Talley, 2006;Kessler, 2006] where phytoplankton in the equatorial undercurrentdisplay only weak seasonality [Pennington et al.,2006]. The EEP is generally a region of weak seasonality[Chavez and Toggweiler, 1995; Loubere, 1998; Loubereand Fariduddin, 1999] and high-nitrate low-chlorophyllconcentration [Behrenfeld and Kolber, 1999; Penningtonet al., 2006]. This means that upwelled nutrients are neverfully utilized by the plankton [Chavez and Barber, 1991]because of iron and silica limitation [Dugdale et al., 1995,2002; Dugdale and Wilkerson, 1998]. The Costa Rica Domeis an oceanic upwelling center in the EEP along the coasts ofNicaragua and Costa Rica where the thermocline approachesvery near the sea surface [Fiedler and Talley, 2006].[17] Because the EEP is a major upwelling zone, allsurface ocean parameters we consider herein have steepgradients across the region. This makes the area an idealstudy site for the effects of environmental factors onsedimentary CaCO 3 dissolution proxies using tropicalplanktonic foraminifers. However, precisely because it isan upwelling zone, all surface ocean parameters in thisregion tend to covary (Table 1) which makes distinguishingthe effect of one parameter (e.g., temperature) from another(e.g., [NO 3 ]) challenging.3. Methods3.1. Samples[18] We used whole tests from P. obliquiloculata andN. dutertrei for SNSW because those two species are mostcommonly used in dissolution work [e.g., Broecker andClark, 2001a, 2001b; Dekens et al., 2002; Mekik andFrançois, 2006; Naik and Naidu, 2007]. Because there isa steep gradient to temperature between 50 and 150 m waterdepth in the EEP, it is important to identify more precisedepth habitats for each of our species. All of our species arethermocline dwellers [Bé, 1960; Hilbrecht, 1996; Anand etal., 2003; Farmer et al., 2007]; however, P. obliquiloculataprefer living at 50 m water depth [Farmer et al., 2007] andMekik et al. [2007a] found the best relationship betweenMg/Ca in shells of this species and water temperatures at50 m. G. menardii prefer 75 m water depth [Farmer et al.,2007] while N. dutertrei are known to live in the deepchlorophyll maximum (DCM) [Fairbanks et al., 1982;Fairbanks and Wiebe, 1980; Loubere, 2001]. We usedLoubere’s [2001] habitat depth estimates for N. dutertreisbased on isotope equilibrium depths calculated using acombination of d 13 C and d 18 O from N. dutertrei shells inEEP core tops. For samples beyond Loubere’s [2001] dataset (those from the OJP and some samples from the EEP),we used the average value of the environmental parameterof interest between 50 and 75 m because this is the meanvalue of habitat depth in Loubere’s [2001] work and thisdepth is also in keeping with independent DCM depthestimates for the EEP [Fiedler and Talley, 2006; Kessler,2006; Pennington et al., 2006].[19] We picked N. dutertrei shells from the 355–415 mmsize range as described by Broecker and Clark [2001a,2001b], and P. obliquiloculata’s from the 420–520 mm sizerange because smaller P. obliquiloculata shells are notabundant in our samples from the EEP. Many studies haveillustrated that using larger foraminifers improves analyticalaccuracy because larger foraminifer size minimizes ontogeneticeffects [Kroon and Darling, 1995] and provides moreconsistent results among samples [Oppo and Fairbanks,1989]. Even in Mg/Ca work on planktonic foraminifershells, Elderfield et al. [2002] established that geochemicaldata from larger foraminifers yield results which are moreconsistent with temperatures in foraminifer habitat waters.3of15

PA1216MEKIK AND RATERINK: CALCITE DISSOLUTION IN THE DEEP SEAPA1216Table 1. Correlations Between Variables in the EEP aWaterCalciteTemperature Nitrate AOU2CO 3R 2 Depth, m2DCO 3 Dissolved, % 50 m 75 m DCM 50 m 75 m DCM 50 m 75 m 100 m 50 m 75 m 100 mWater depth 1 0.33 0.04 0.09 0.07 0 0.06 0.06 0.01 0.12 0.1 0.05 0.08 0.06 0.01DCO 3 0.33 1 0.18 0.01 0.01 0.05 0.02 0 0.06 0.01 0 0.2 0.02 0.03 0.11Percent dissolved 0.04 0.18 1 0.18 0.24 0.3 0.28 0.36 0.52 0.21 0.28 0.38 0.25 0.36 0.48T 50 m 0.09 0.01 0.18 1 0.79 0.5 0.6 0.44 0.43 0.82 0.64 0.54 0.85 0.68 0.58T 75 m 0.07 0.01 0.29 0.79 1 0.61 0.75 0.72 0.75 0.82 0.83 0.78 0.86 0.88 0.81T DCM 0 0.05 0.3 0.5 0.61 1 0.58 0.54 0.78 0.5 0.52 0.84 0.56 0.58 0.93Nitrate 50 m 0.06 0.02 0.28 0.6 0.75 0.58 1 0.82 0.78 0.75 0.76 0.73 0.86 0.86 0.84Nitrate 75 m 0.06 0 0.36 0.44 0.72 0.54 0.82 1 0.92 0.66 0.87 0.85 0.67 0.88 0.83Nitrate DCM 0.01 0.06 0.52 0.43 0.75 0.78 0.78 0.92 1 0.58 0.81 0.9 0.6 0.84 0.9AOU 50 m 0.12 0.1 0.21 0.82 0.82 0.5 0.75 0.66 0.58 1 0.83 0.71 0.89 0.83 0.68AOU 75 m 0.1 0 0.28 0.64 0.83 0.52 0.76 0.87 0.81 0.83 1 0.87 0.8 0.93 0.79AOU DCM 0.05 0.02 0.38 0.54 0.78 0.84 0.73 0.85 0.9 0.71 0.87 1 0.68 0.83 0.912CO 3 50 m 0.08 0.02 0.25 0.85 0.86 0.56 0.86 0.67 0.6 0.89 0.8 0.68 1 0.87 0.752CO 3 75 m 0.06 0.03 0.36 0.68 0.88 0.58 0.86 0.88 0.84 0.83 0.93 0.83 0.87 1 0.87DCM 0.01 0.11 0.48 0.58 0.81 0.93 0.84 0.83 0.9 0.68 0.79 0.91 0.75 0.87 1CO 32a T is temperature; DCM stands for deep chlorophyll maximum.Furthermore, J. Bijma’s (personal communication, 2006)unpublished SNSW data from Globigerinoides sacculifertests also supports the assertion that analytical accuracyimproves with increasing foraminifer size.[20] We ascertain that our core tops are Holocene inseveral ways. First, Loubere [2001] provided d 13 C andd 18 O measurements from N. dutertrei shells in core topsfrom the EEP which are consistent with Holocene d 13 C andd 18 O values. Some of the core tops we are using hereinoverlap with a subset of Loubere’s [2001] samples, and thegeographic distribution of his core tops is broad enough toallow for a regional assessment of age distribution in theEEP. In addition to Loubere’s [2001] data set, Mekik et al.[2007b] generated d 18 O data from foraminifers in samplesfrom very deep core tops where chemical erosion may haveobliterated Holocene sediments and they excluded sampleswith values inconsistent with those for the Holocene. Wealso excluded those samples herein. Second, Mekik et al.’s[2007b] EEP rain ratio maps generated from a subset ofsamples used herein fit well with chlorophyll-based estimatesof Recent surface ocean productivity in the EEP [afterBehrenfeld and Falkowski, 1997]. Last, Mekik et al.’s[2007a] Mg/Ca data from planktonic foraminifers fromthe same core tops as those used herein show consistentpatterns with Recent sea surface temperatures in the habitatwaters of each species.3.2. Sample Preparation[21] We followed methods outlined by Mekik et al. [2002]for generating MFI data, and used procedures described byLohmann [1995] and Broecker and Clark [2001a] forSNSW measurements with three modifications in order toimprove data quality. First, foraminifers were picked individuallywithin given size ranges instead of trapping foraminifersbetween two sieves; and all the picking was doneby the same person (F. Mekik). This ensures that sizemeasurements for each foraminifer are made on the shortestdiameter and improves the consistency of foraminifer sizewithin each range. Second, a wet picking technique wasused because wet foraminifers are more transparent. Thisfacilitates picking the cleanest specimens because using anultrasonicator or other cleaning methods on foraminifersgenerally damages shells and distorts weight data. Third, weweighed two separately picked, size-normalized populationsfor each species from each sample in order to comparereplicate measurements of species-specific mean weightfrom each sediment sample. Replicate mean weight measurementsfor all samples are not available, however, becauseof the low abundance of clean foraminifers in given sizeranges in some samples.[22] We aimed for 50 or more whole shells for weighingfrom each sediment sample. However, in some samples,especially those from high-dissolution areas, we had to baseour mean weight on fewer foraminifers. Our data set forMFI (Figure 1) is larger than that for SNSW becausefragments of G. menardii are far more abundant than cleanwhole foraminifer shells in tight size ranges.3.3. Analytical Method[23] Mekik et al. [2002] estimated MFI’s error margin at10–15%. It is customary to count 300 or more G. menardiiwhole shells and fragments per sample to obtain statisticallyrobust results. Counts in some high-dissolution samples fellbelow 300 because of lack of shell matter.[24] There are three main sources of uncertainty in SNSWmeasurements: (1) error margin of the balance; (2) reproducibilityof mean weight measurements; and (3) variationin shell weight within a given size range. We used a Mettlermicrogram scale at Grand Valley State University whoseerror margin is ±5 mg. To test for reproducibility of meanweight measurements, we made replicate weight measurementsfrom a separately picked, second population offoraminifers from the same species and in the same sizefractions from each sediment sample where we had asufficient number of foraminifers available. The reproducibilitybetween two separately picked replicate mean weightmeasurements is high (R 2 = 0.97). In the following discussion8 represents the difference between two replicatemean weight measurements for the same species in the samesediment sample. Mean 8 for P. obliquiloculata weight4of15

PA1216MEKIK AND RATERINK: CALCITE DISSOLUTION IN THE DEEP SEAPA1216Figure 1. (a) General locations for samples used herein are shown as open and shaded diamonds. OJP isOntong Java Plateau; EEP is eastern equatorial Pacific. Dashed lines represent temperature contours for75 m water depth after Locarnini et al. [2006]. (b) Core top sample distribution in the eastern equatorialPacific. MFI data are available from all samples. Shaded diamonds show samples from which N. dutertreiSNSW were generated, and dots show samples from which P. obliquiloculata SNSW data weregenerated. Dashed lines represent seafloor bathymetry.measurements is 3.74 mg, and the standard deviation in 8for P. obliquiloculata weight is 3.02 mg. The ratio of mean 8to average P. obliquiloculata mean weight from all samples(65 mg) is 5.8%. For N. dutertrei weight data, mean 8 is2.18 mg, and its standard deviation is 2.02 mg. The ratio formean 8 to average N. dutertrei mean weight from allsamples (37.5 mg) is also 5.8%.[25] Following the work of others [Rosenthal et al., 2000;Barker et al., 2004], we examined single-foraminiferweights within three size ranges for P. obliquiloculatawhole shells in a sample from the OJP, ERDC 89. Thevariation coefficient for each size range is listed as apercentage (Table 2) and it is the ratio of the standarddeviation for all weight measurements within the size rangeto the mean weight of foraminifers in that size range. Notethat the variation coefficient for P. obliquiloculata weight inthe 355–420 mm size range is significantly higher than thatfor the two larger and wider size ranges. These resultsconfirm our choice for using larger P. obliquiloculataspecimens because weight variation within a size rangeseems to be lower among larger foraminifers. We did notgenerate single-foraminifer weight data for N. dutetreibecause an average N. dutertrei shell in the 355–420 mmsize fraction weighs 38 mg. The uncertainty of the balanceat ±5 mg would substantially obscure the variation in ourweight measurements of individual N. dutertrei shells.[26] Data for [CO 3 2 ] at 50, 75 and 100 m water depth arefrom Archer [1996, personal communication, 2001]. Otherenvironmental data (temperature, [NO 3 ] and AOU) at 50,75 and 100 m are from NOAA’s World Ocean Atlases (seeTable 2. Variation Coefficients in Single P. obliquiloculata ShellWeights in Sample ERDC 89 on the Ontong Java Plateau355–420 mm 420–520 mm 520–620 mmMean weight, mg 61.1 90.8 122.4Standard deviation 7.6 8.4 10.5Variation coefficient, % 12.4 9.3 8.65of15

PA1216MEKIK AND RATERINK: CALCITE DISSOLUTION IN THE DEEP SEAPA1216Locarnini et al. [2006] for temperature and Garcia et al.[2006a, 2006b] for nutrients).4. Results[27] Our sediment samples include core tops along (1) adepth transect on the OJP where surface ocean parametersare mostly invariable allowing us to isolate the dissolutionsignal in our proxies; and, for comparison, (2) a large groupof core top samples (more than 100) from the EEP (1808–4440 m depth) where surface ocean parameters are stronglyvariable (Figure 1). All samples are from gravity cores.Listings for all data used herein are available as auxiliarymaterial. 14.1. Calcite Dissolution Proxies on the Ontong JavaPlateau[28] First we compare MFI and SNSW in a subset of MFI’scalibration samples in core tops from the OJP. These samplescompose a depth transect (1900–4441 m) beneath surfacewaters with fairly similar environmental parameters. Thisallows us to isolate the CaCO 3 dissolution signal in ourproxies.2[29] On OJP, MFI has a robust relationship with DCO 3(R 2 = 0.92, Figure 2a), model-derived percent CaCO 3dissolved (R 2 = 0.88, Figure 2d), and Mg/Ca fromP. obliquiloculata shells (R 2 = 0.94, Figure 2g). However,neither P. obliquiloculata nor N. dutertrei shell weight2correlates very well with DCO 3 (Figures 2b and 2c),MFI-based percent CaCO 3 dissolved (Figures 2e and 2f)or Mg/Ca from P. obliquiloculata shells (Figures 2h and 2i).[30] Though P. obliquiloculata shell weight drops withincreasing dissolution in OJP samples (Figures 2b and 2e),we do not see a complementing relationship between MFIand SNSW there. Instead, increasing fragmentation inG. menardii shells appears to be happening under the same2conditions of bottom water CO 3 undersaturation as wholeshell weight loss in P. obliquiloculatas (Figures 2a, 2b, 2d,and 2e). In addition, Mg/Ca in P. obliquiloculata shells alsohas a weak relationship with P. obliquiloculata SNSW(Figure 2h).[31] N. dutertrei shells are significantly less abundant andlighter in OJP samples (OJP average weight is 27 mg) whencompared to their counterparts in the EEP (EEP averageweight is 39 mg). Broecker and Clark [2001a, 2001b] alsolist lighter N. dutertrei weights (20–36 mg) in their samplesfrom the OJP within the same size fraction we used here.4.2. Calcite Dissolution Proxies in the EasternEquatorial Pacific4.2.1. Globorotalia menardii Fragmentation Index[32] The fragmentation trend of G. menardii (MFI) hasno clear mathematical relationship with any surface ocean2parameter or DCO 3 of bottom waters in the EEP (Figure 3).At the same temperature at 75 m water depth (16°C) wesee a wide range to MFI (0.4–1). Similarly, samples under awide range of temperature (15°–22°C) yield more or lessconstant MFI values (1) (Figure 3). We find similar1 Auxiliary materials are available at to that of MFI versus temperature when we plotMFI against surface ocean [CO 23 ], [NO 3 ] and AOU(Figure 3). The poor relationship between MFI and bottom2water DCO 3 in the EEP is not unexpected [Mekik et al.,2002]. Unlike on the OJP where there is little variation inseabed organic carbon flux, respiration of carbon in sedimentpore waters drives additional CaCO 3 dissolutionwithin sediments in the EEP [Mekik et al., 2002, 2007b].4.2.2. Size-Normalized Shell Weight[33] P. obliquiloculata shell weight correlates well withMFI-based percent dissolved in the EEP (R 2 = 0.79)although it has no correlation with DCO 23 . This is unexpectedbecause we were not able to find a good mathematicalrelationship between P. obliquiloculata shell weight andMFI-based percent dissolved or Mg/Ca in P. obliquiloculatashells in samples from our depth transect on the OJP wheretemperature and other surface ocean parameters are mostlyunchanging (Figures 2b, 2e, and 2h). It appears that in theEEP, P. obliquiloculata shell weight is influenced by bothrespiratory CaCO 3 dissolution within the sediments and, to alesser extent, environmental parameters in the surface ocean.Also, P. obliquiloculata SNSW has somewhat variable butweaker relationships with all four environmental parametersat 50 m water depth (R 2 = 0.46–0.77) (Figure 4). We areusing exponential relationships because they provide thebest fit with our data.[34] N. dutertrei SNSW in core tops from the EEP(Figure 5) show no significant correlation with any dissolutionor environmental parameter. However, we see a moderatelystrong correlation between MFI-based percentCaCO 3 dissolved and N. dutertrei SNSW (Figure 5b) ifequatorial samples are removed. We also observe a correlationbetween [CO 23 ] and/or [NO 3 ] at the DCM andN. dutertrei SNSW, again if equatorial samples are removed(Figures 5d and 5e). Furthermore, N. dutertrei shell weightin equatorial samples seem to be heavier than what thetrend seen in other samples indicates (Figures 5c–5f); andN. dutertrei weights from just north of the equator areheavier than those from immediately south of the equator.5. Discussion5.1. Comparison of Dissolution Proxies on the OntongJava Plateau[35] Both our SNSW data using P. obliquiloculata shellsand those of Broecker and Clark [2001a] using smallerindividuals of the same species have weak correlations with2DCO 3 on the OJP (Figure 2b). This is different than therobust relationship between P. obliquiloculata SNSW and2bottom water CO 3 undersaturation [Broecker and Clark,2001a].[36] The discrepancy lies in the different terms used to2describe undersaturation of CO 3 in bottom waters.Broecker and Clark [2001a] use [CO 23 ]*, which theydefine as a depth normalized bottom water [CO 23 ] indicatorbased on [CO 23 ] values extrapolated from GEOSECS. The2DCO 3 values given by Archer [1996] are also derivedfrom GEOSECS, but these are based on the empiricalrelationship between [CO 2 3 ] and temperature, salinity, O 2and nutrients, and CaCO 3 solubility formulations. We use6of15

PA1216MEKIK AND RATERINK: CALCITE DISSOLUTION IN THE DEEP SEAPA1216Figure 2. All samples shown here are from OJP. (a) MFI versus DCO 2 3 .(b)P. obliquiloculata SNSWversus DCO 23 . Stars show SNSW data (355–415 mm) from Broecker and Clark [2001a]. Diamondsshow data from this study (420–520 mm). (c) N. dutertrei SNSW versus DCO 23 . Stars show SNSW datafrom Broecker and Clark [2001a, 2001b]. Diamonds show data from this study. Both are from the355–420 mm size fraction. (d) MFI versus model-derived estimates of percent calcite dissolved. (e)P. obliquiloculata SNSW versus MFI-based percent calcite dissolved estimates in MFI’s calibrationsamples. (f) N. dutertrei SNSW versus MFI-based percent calcite dissolved estimates. (g) MFI versusMg/Ca in P. obliquiloculata shells. (h) P. obliquiloculata SNSW versus Mg/Ca in P. obliquiloculata.(i) N. dutertrei SNSW versus Mg/Ca in P. obliquiloculata.2Archer’s [1996] DCO 3 data herein for comparing MFIwith SNSW in order to maintain consistency with formerwork on MFI [Mekik et al., 2002; Mekik and François,2006; Mekik et al., 2007a, 2007b], and because MFI dataare not available for the sample set used by Broecker andClark [2001a].5.2. Postdepostional Calcite Dissolution andEnvironmental Parameters in Foraminifer HabitatWaters[37] It is difficult to isolate the influence of a singlespecific surface ocean parameter on CaCO 3 dissolutionproxies in the EEP for two reasons. First, all our core tops7of15

PA1216MEKIK AND RATERINK: CALCITE DISSOLUTION IN THE DEEP SEAPA12162Figure 3. MFI versus DCO 3 and environmental parameters at 75 m water depth in core top samples2from the eastern equatorial Pacific. Shaded line in first plot showing MFI versus DCO 3 represents thecorrelation line between these same parameters in core tops from the OJP (Figure 2a).are from regions where bottom waters are undersaturated2 2with respect to CO 3 (DCO 3 ranges between 0.42 and41.8 mmol/kg among locations of our samples). Thus wedo not have samples that experienced no dissolution. Evenif we had very shallow samples (

PA1216MEKIK AND RATERINK: CALCITE DISSOLUTION IN THE DEEP SEAPA12162Figure 4. P. obliquiloculata SNSW data versus (a) DCO 3 and environmental parameters (b) MFIpercent dissolved calcite, (c) water temperature, (d) carbonate ion and (e) nitrate concentrations, and(f) apparent oxygen utilization at 50 m water depth in core top samples from the eastern equatorial Pacific.In Figure 4b, color coding indicates [CO 23 ] at 50 m water depth in the EEP. In Figures 4c–4f, colorcoding indicates three dissolution brackets estimated with MFI.9of15

PA1216MEKIK AND RATERINK: CALCITE DISSOLUTION IN THE DEEP SEAPA12162Figure 5. N. dutertrei SNSW data versus (a) DCO 3 and environmental parameters (b) MFI-basedpercent calcite dissolved, (c) water temperature, (d) carbonate ion and (e) nitrate concentrations, and(f) apparent oxygen utilization at the deep chlorophyll maximum (N. dutertrei habitat depths fromLoubere [2001]) in core top samples from the eastern equatorial Pacific. Yellow triangles show samplesfalling in the region between the equator and 5°S, and yellow circles show samples from the regionbetween the equator and 5°N.10 of 15

PA1216MEKIK AND RATERINK: CALCITE DISSOLUTION IN THE DEEP SEAPA1216should not be restricted to O. universa and Globigerinoidessacculifer [Bijma et al., 2002]. Naik and Naidu [2007]provided evidence in core tops from the western tropicalIndian Ocean supporting the strong effect of [CO 23 ] incalcification waters on SNSW of both N. dutertrei andP. obliquiloculata while Barker and Elderfield [2002]demonstrated the same effect in down core work. Evenstudies on shell chemistry, not just SNSW, show that the[CO 23 ] of habitat waters may bias results (e.g., d 18 O andd 13 C[Spero et al., 1997] and U/Ca [Russell et al., 2004]).Although our data set does not allow us to rule out thepotential effect of the other three environmental parameterson SNSW (because all surface ocean parameters covary inthis region), previous studies point to [CO 23 ] in calcificationwaters as the most influential environmental parameteron shell weight. So, we will focus our discussion around theeffect of [CO 23 ] in habitat waters on SNSW in the EEP.[39] In order to examine the effect of [CO 2 3 ] of habitatwaters on the dissolution trend seen in P. obliquiloculatashell weight, we grouped our samples into three rangesof [CO 23 ] at 50 m water depth (220 mmol/kg) (Figure 4b). If calcite dissolution werethe only influence on P. obliquiloculata shellweightinthe EEP, we would expect the distribution of [CO 23 ]valuesinFigure 4b to be overlapping and random. Instead, SNSW ofP. obliquiloculata (Figure 4) seems to respond both topostdepositional CaCO 3 dissolution in the sediments (R 2 =0.79; estimated with MFI) and, most likely, [CO 2 3 ]ofwatersat 50 m in the EEP. Our results are supported by Naik andNaidu’s [2007] findings of the strong influence of [CO 23 ]incalcification waters on P. obliquiloculata SNSW.[40] Likewise, we grouped samples into three categoriesbased on the extent of dissolution each experienced (60%). The width of each dissolution category(15%) in Figures 4c–4f is within MFI’s errormargin. Again, both dissolution and [CO 23 ] in ambientwaters appear to influence P. obliquiloculata shell weight(Figure 4d), but there is also a systematic decrease in thesensitivity of P. obliquiloculata SNSW to [CO 23 ] in sampleswhich have experienced high dissolution (green dots).[41] We performed multiple linear regression analysis tofurther explore the effect of postdepositional calcite dissolutionand [CO 23 ] at 50 m water depth on P. obliquiloculataSNSW. In our analysis P. obliquiloculata SNSW is thedependent variable and MFI-based percent CaCO 3 dissolvedand [CO 23 ] at 50 m are the independent variables.We used the resulting multiple linear regression equation(Figure 6) to estimate P. obliquiloculata SNSW from MFIbasedpercent CaCO 3 dissolved and [CO 23 ] at 50 m. waterdepth as input parameters for each sample. We find a highcorrelation between measured P. obliquiloculata weightsand those calculated with the regression equation (Figure 6).This suggests that 83% of the variation in P. obliquiloculataSNSW in our samples may be explained by the effects ofboth postdepositional CaCO 3 dissolution and [CO 23 ]at50mwater depth (Figure 6).5.3. Geographic Controls on Dissolution Proxies[42] All environmental variables in the EEP covary(Table 1), but we believe the environmental factor mostFigure 6. Multiple linear regression equation betweenMg/Ca in P. obliquiloculata and MFI-based percent calcitedissolved and [CO 23 ] at 50 m in samples from the EEP.likely affecting N. dutertrei shell weight in our EEP samplesis [CO 23 ] at DCM depths (Figure 5d) because of stronglaboratory evidence supporting the effect of [CO 23 ] inambient waters on shell weight [Bijma et al., 1999, 2002]and core top work showing the influence of surface ocean[CO 2 3 ]onN. dutertrei SNSW [Naik and Naidu, 2007].[CO 23 ] at the DCM seems to increase from south to northacross the EEP with the exception of relatively low [CO 2 3 ]in the equatorial region.[43] Within the EEP, areas beneath the North EquatorialCounter Current and in the southwest beyond the upwellingzone contain the heaviest N. dutertrei shells within ourgiven size ranges (Figure 7). Both of these areas have verylow chlorophyll concentrations (low productivity) in surfacewaters. By contrast, lighter N. dutertrei shells are foundbeneath upwelling regions on the equator by the SouthEquatorial Current (Figure 7) and at the Costa Rica Dome[Fiedler and Talley, 2006], where productivity is high.[44] The geographic pattern of SNSW for N. dutertreipopulations in the EEP may indicate the presence of‘‘cryptic species’’ of N. dutertrei, which may prefer specificenvironmental factors [Darling et al., 1996; Huber et al.,1997; Darling et al., 1999; Kucera and Darling, 2002;Darling et al., 2003]. Cryptic species are distinct genotypeswithin a morphospecies that can be difficult to identify withmorphological features alone. The distribution of N. dutertrei11 of 15

PA1216MEKIK AND RATERINK: CALCITE DISSOLUTION IN THE DEEP SEAPA1216Figure 7. Distribution of N. dutertrei SNSW in the eastern equatorial Pacific. Abbreviations areNEqCC, North Equatorial Counter Current; SEqC, South Equatorial Current; and CRD, Costa RicaDome. Contours represent N. dutertrei shell weight in mg. The image is overlain on the chlorophyll-basedsurface ocean productivity map of Behrenfeld and Falkowski [1997].SNSW could reflect intraspecific ecophenotypic variation,multiple genotypes or both, particularly where the SNSWdoes not seem to be influenced by [CO 23 ] (Figure 5d). Wewould need DNA data from our N. dutertrei specimens todiscern among these possibilities, but this data is not available.We note, though, that Kucera and Darling [2002], usingDNA data, describe three distinct genotypes for N. dutertreibut only one for P. obliquioculata and one for G. menardii.This finding may explain the regional distribution ofN. dutertrei shell weight as evidence of multiple populationsof the N. dutertrei morphotype in the EEP distinguished bytheir shell thickness.[45] Schmidt et al. [2003, 2004] reported that planktonicforaminifer size relates to latitude. They attribute thischange to surface water temperature, because higher temperaturespromote growth in foraminifers. Schmidt et al.[2004] also conclude that planktonic foraminifer assemblagestend to be smaller in upwelling regions. Although theydo not present SNSW data, it is possible that the SNSWdistribution of N. dutertrei across the EEP (Figure 7) reflectsvariations in temperature and upwelling.[46] Furthermore, the SNSW of N. dutertrei also variesbetween the OJP and EEP, with much lighter specimens onthe OJP. Broecker and Clark [2001a] note that N. dutertreishells from the Atlantic Ocean are heavier than those fromthe Indian and Pacific Oceans. Thus there appears to bemultiple populations of the N. dutertrei morphotype both inthe EEP and across the equatorial Pacific between the EEPand OJP.[47] One last geographic control on dissolution proxiesstems from the absence of G. menardiis in Atlantic sedimentsfrom the Last Glacial Maximum (LGM). This limitsMFI’s applicability in down core work in Atlantic cores.5.4. Other Complicating Factors[48] A factor often ignored in studies using foraminiferSNSW is that significant shell loss may occur as foraminifertests settle though the water column; however, this has notbeen well established for many foraminifer species becauseof the scarcity of sediment trap data. Schiebel [2002]estimated that only 25% of initially produced foraminifershell material settles to the bottom. In a more recent study,Schiebel et al. [2007] illustrated that Globigerina bulloidesand Globigerinita glutinata lose an average of 19% of theiroriginal shell weight while settling through the twilightzone, 100–1000 m water depth. They detected no averagetest weight loss below the twilight zone, and report thatforaminifer shells may even gain weight there. On the otherhand, most foraminifer tests experiencing dissolutionthrough the water column likely belong to juveniles becausemost foraminifers in sediments are gametogenic and probablysank to the seabed rapidly after death (J. Bijma,personal communication, 2006). Thus dissolution of adulttests during sinking seems unlikely. Moreover, fragmentationof G. menardii shells in the water column has not beenreported. G. menardiis are known to be somewhat resistantto dissolution even in sediments [Berger, 1968, 1970;Thompson and Saito, 1974] and are used often to estimate12 of 15

PA1216MEKIK AND RATERINK: CALCITE DISSOLUTION IN THE DEEP SEAPA1216deep-sea CaCO 3 dissolution [Oba, 1969; Ku and Oba,1978; Peterson and Prell, 1985a, 1985b; Mekik et al.,2002; Mekik and François, 2006].6. Conclusions[49] We explored the significance of environmental influenceson two deep-sea CaCO 3 dissolution proxies in a largenumber of core top samples from the tropical Pacific: sizenormalizedforaminifer shell weight and the G. menardiifragmentation index. We find that SNSW and MFI do notcomplement each other and, instead, trace dissolutionconcurrently in core tops from both the OJP and theEEP. The dissolution signal in P. obliquiloculata SNSWis weak in samples from the OJP where surface oceanparameters are mostly constant. Conversely, SNSW ofP. obliquiloculata shells in samples from the EEP carriesa strong dissolution signal. Though we cannot isolate whichenvironmental parameter is affecting P. obliquiloculatashell weight in the EEP with our data, we are able to showthat P. obliquiloculata SNSW responds to both postdepositionalshell dissolution and surface ocean parameters there.N. dutertrei SNSW, on the other hand, shows a distinctlatitudinal pattern in the EEP in keeping with regional highproductivityzones and current systems suggesting thatthere may be intraspecific ecophenotypic variations in thismorphotype in the EEP which may be reflected in itsSNSW.[50] Our study is not the first to find an environmentalinfluence of foraminifers’ calcification waters on theirSNSW. This is the first study, however, where environmentaleffects on SNSW are documented with core topsediment samples in the EEP where calcite dissolution is2driven by both bottom water DCO 3 and organic carbondegradation in sediment pore waters. This is important forpaleoceanographic work because calibration equationsfrom core top sediment samples are often used for downcore applications and because culture experiments andsediment trap data are few. Furthermore, it is not possibleto reliably study postdepositional shell dissolution in laboratorywork and with sediment traps because the effect oforganic carbon degradation in sedimentary pore waters oncalcite dissolution is difficult to mimic in a laboratorysetting.[51] Although our results cannot offer a clear mechanisticexplanation for the variation in shell weight in the EEP andfurther work with laboratory cultures is required to accomplishthis, the correlations we demonstrate between SNSWand both MFI-based percent dissolved values and surfaceocean parameters, particularly [CO 23 ], cannot be ignored indown core applications. Our findings suggest that cautionmust be used when making paleoceanographic inferencesfrom SNSW variations in the paleorecord.[52] Finally, we show that MFI-based percent CaCO 3dissolved estimates are mostly insensitive to surface oceanenvironmental parameters in G. menardii’s calcificationwaters in the EEP. Mekik and François [2006] showedlinear decreases in Mg/Ca and Mg/Sr in P. obliquiloculataand G. menardii shells with increasing dissolution estimatedusing MFI. Mekik et al. [2002, 2007a, 2007b] demonstratedMFI’s applicability outside its calibration area in core topsamples from the EEP upwelling region. Loubere et al.[2004] and Richaud et al. [2007] showed MFI’s applicabilityin down core work in reconstructing paleocalcite fluxes.With all of these qualities, MFI seems to approach ourdefinition of the ideal CaCO 3 dissolution proxy described inthe introduction of this paper with three caveats: (1) abiological/ecological bias in MFI remains to be explored;(2) MFI’s range of percent dissolved is still limited to 25–76%; and (3) G. menardiis are absent in Atlantic sedimentsfrom the LGM. Thus the ideal CaCO 3 dissolution proxy isstill elusive.[53] Acknowledgments. This manuscript benefited substantiallyfrom many fruitful discussions with Paul Loubere and Roger François.Constructive and thoughtful comments by Jerry Dickens and three anonymousreviewers much improved our manuscript. We gratefully acknowledgethe curators and repositories that provided sediment samples and helpin selecting cores for this work (June Padman, Oregon State University;Larry Peterson, RSMAS; Rusty Lotti-Bond, Lamont-Doherty Earth Observatory;Warren Smith, Scripps Institution of Oceanography; and curators atthe University of Hawaii). Thanks also go to the National ScienceFoundation for the support it provides to those repositories. This studywas supported in full by grant OCE0326686 from the National ScienceFoundation.ReferencesAnand, P., H. Elderfield, and M. H. Conte(2003), Calibration of Mg/Ca thermometry inplanktonic foraminifera from a sediment traptime series, Paleoceanography, 18(2), 1050,doi:10.1029/2002PA000846.Archer, D. E. (1996), An atlas of the distributionof calcium calcite in sediments of the deep sea,Global Biogeochem. Cycles, 10, 159–174.Archer, D., and E. Maier-Reimer (1994), Effectof deep sea sedimentary calcite preservation onatmospheric CO 2 concentration, Nature, 367,260–264.Archer, D., A. Winguth, D. Lea, and N. Mahowald(2000), What caused the glacial/interglacialatmospheric pCO 2 cycles?, Rev. Geophys.,38, 159 – 189.Archer, D. E., J. L. Morford, and S. Emerson(2002), A model of suboxic sedimentary diagenesissuitable for automatic tuning and griddedglobal domains, Global Biogeochem. Cycles,16(1), 1017, doi:10.1029/2000GB001288.Arrhenius, G. (1952), Sediment cores from theeast Pacific, Rep. Swed. Deep Sea Exped.1947–1948, 5, 1 –228.Barker, S., and H. Elderfield (2002), Foraminiferalcalcification response to glacial-interglacialchanges in atmospheric CO 2 , Science, 297,833–836.Barker, S., K. Thorsten, and H. Elderfield(2004), Temporal changes in North Atlanticcirculation constrained by planktonic foraminiferalshell weights, Paleoceanography, 19,PA3008, doi:10.1029/2004PA001004.Bé, A. W. H. (1960), Ecology of Recentplanktonic foraminifera: part 2—Bathymetricand seasonal distributions in the SargassoSea off Bermuda, Micropaleontology, 6,373–392.Behrenfeld, M., and P. Falkowski (1997), Photosyntheticrates derived from satellite-basedchlorophyll concentration, Limnol. Oceanogr.,42, 1 – 20.Behrenfeld, M., and Z. S. Kolber (1999), Widespreadiron limitation of phyroplankton in theSouthern Ocean, Science, 283, 840–843.Berger, W. (1968), Planktonic foraminifera: Selectivesolution and paleoclimatic interpretation,Deep Sea Res. Oceanogr. Abstr., 15, 31–43.Berger, W. (1970), Planktonic foraminifera: Selectivesolution and the lysocline, Mar. Geol.,8, 111–138.Berger, W. (1973), Deep sea carbonates; Pleistocenedissolution cycles, J. Foraminiferal Res.,3, 187–195.13 of 15

PA1216MEKIK AND RATERINK: CALCITE DISSOLUTION IN THE DEEP SEAPA1216Berger, W. (1989), Global maps of ocean productivity,in Productivity of the Ocean: Presentand Past, edited by W. H. Berger, V. S. Smetacek,and G. Wefer, pp. 429 –455, John Wiley,New York.Berger, W., K. Fischer, C. Cai, and G. Wu(1987), Organic productivity and organic carbonflux, I, in Overview and Maps of PrimaryProduction and Export Production, Rep. 87-30, pp. 1–45, Scripps Inst. of Oceanogr., Univ.of Calif., La Jolla.Bijma, J., H. Spero, and D. W. Lea (1999), Reassessingforaminiferal stable isotope geochemistry:Impact of the oceanic carbonatesystem (experimental results), in Uses ofProxies in Paleoceanography: Examples fromthe South Atlantic, editedbyG.FischerandG. Wefer, pp. 489–512, Springer, New York.Bijma, J., B. Honisch, and R. E. Zeebe (2002),The impact of the ocean carbonate chemistryon living foraminiferal shell weight: Commenton ‘‘Carbonate ion concentration in glacialagedeep waters of the Caribbean Sea’’ byW. S. Broecker and E. Clark, Geochem. Geophys.Geosyst., 3(11), 1064, doi:10.1029/2002GC000388.Broecker, W. S. (1971), A kinetic model for thechemical composition of sea water, Quat. Res.,1, 188–207.Broecker, W. (1982), Ocean chemistry duringglacial time, Geochim. Cosmochim. Acta, 46,1689–1705.Broecker, W. S., and E. Clark (2001a), An evaluationof Lohmann’s foraminifera weight dissolutionindex, Paleoceanography, 16, 431–434.Broecker, W. S., and E. Clark (2001b), Glacialto-Holoceneredistribution of carbonate ion inthe deep sea, Science, 294, 2152–2155.Broecker, W. S., and E. Clark (2003), Glacialagedeep sea carbonate ion concentrations,Geochem. Geophys. Geosyst., 4(6), 1047,doi:10.1029/2003GC000506.Brown, S. J., and H. Elderfield (1996), Variationsin Mg/Ca and Sr/Ca ratios of planktonic foraminiferacaused by postdepositional dissolution:Evidence of shallow Mg-dependentdissolution, Paleoceanography, 11, 543–551.Chavez, F. P., and R. T. Barber (1987), An estimateof new production in the equatorial Pacific,Deep Sea Res., Part A, 34, 1229–1243.Chavez, F. P., and R. T. Barber (1991), The GalapagosIslands and their relation to oceanographicprocesses in the tropical Pacific, inGalapagos Marine Invertebrates, editedbyM. J. James, pp. 9–33, Plenum, New York.Chavez, F. P., and J. R. Toggweiler (1995), Physicalestimates of global new production: Theupwelling contribution, in UpwellingintheOcean: Modern Processes and Ancient Records,edited by C. P. Summerhayes et al.,pp. 313–320, John Wiley, Chichester, UK.Darling, K. F., D. Kroon, C. M. Wade, and A. J.Leigh Brown (1996), Molecular evolution ofplanktic foraminifera, J. Foraminiferal Res.,26, 324–330.Darling,K.,C.M.Wade,D.Kroon,A.J.L.Brown,and J. Bijma (1999), The diversity and distributionof modern planktonic foraminiferal small subunitribosomal RNA genotypes and their potentialas tracers of present and past ocean circulations,Paleoceanography, 14, 3–12.Darling, K. F., M. Kucera, C. M. Wade, P. vonLangen, and D. Pak (2003), Seasonal distributionof genetic types of planktonic foraminifermorphospecies in the Santa Barbara Channeland its paleoceanographic implications,Paleoceanography, 18(2), 1032, doi:10.1029/2001PA000723.Dekens, P. S., D. W. Lea, D. K. Pak, and H. J.Spero (2002), Core top calibration of Mg/Ca intropical foraminifera: Refining paleo-temperatureestimation, Geochem. Geophys. Geosyst.,3(4), 1022, doi:10.1029/ Villiers, S. (2005), Foraminiferal shell-weightevidence for dissolution in marine sedimentsoverlain by supersaturated bottom waters,Deep Sea Res., Part I, 52, 671–680.Dugdale, R. C., and F. P. Wilkerson (1998), Silicateregulation of new production in theequatorial Pacific upwelling, Nature, 391,270–273.Dugdale, R. C., F. P. Wilkerson, and H. J. Minus(1995), The role of the silicate pump in drivingnew production, Deep Sea Res., Part I, 42,697–719.Dugdale, R. C., A. G. Wischmeyer, F. P.Wilkerson, R. T. Barber, F. Chai, M. S. Jiang,and T. H. Peng (2002), Meridional asymmetryof source nutrients to the equatorial Pacificupwelling ecosystem and its potential impacton ocean-atmosphere CO 2 flux; a data andmodeling approach, Deep Sea Res., Part II,49, 2513–2531.Elderfield, H., and G. Ganssen (2000), Reconstructionof temperature and d 18 O of surfaceocean waters using Mg/Ca of planktonic foraminiferalcalcite, Nature, 405, 442–445.Elderfield, H., M. Vautravers, and M. Cooper(2002), The relationship between shell size andMg/Ca, Sr/Ca, d 18 O, and d 13 C of species of planktonicforaminifera, Geochem. Geophys. Geosyst.,3(8), 1052, doi:10.1029/2001GC000194.Emerson, S., and M. Bender (1981), Carbonfluxes at the sediment-water interface ofthe deep sea: Calcium carbonate preservation,J. Mar. Res., 39, 139–162.Fairbanks, R., and P. Wiebe (1980), Foraminiferaand chlorophyll maximum: Vertical distribution,seasonal succession, and paleoceanographicsignificance, Nature, 1524–1525.Fairbanks, R. G., M. Sverdlove, R. Free, P. H.Wiebe, and A. H. Bé (1982), Vertical distributionand isotopic fractionation of living planktonicforaminifera from the Panama Basin,Nature, 298, 841–844.Farmer, E. C., A. Kaplan, P. B. de Menocal, andJ. Lynch-Stieglitz (2007), Corroborating ecologicaldepth preferences of planktonic foraminiferain the tropical Atlantic with the stableoxygen isotope ratios of core top specimens,Paleoceanography, 22, PA3205, doi:10.1029/2006PA001361.Fiedler, P. C., and L. D. Talley (2006), Hydrographyof the eastern tropical Pacific: A review,Prog. Oceanogr., 69, 143–180.Garcia, H. E., R. A. Locarnini, T. P. Boyer, andJ. I. Antonov (2006a), World Ocean Atlas2005, vol. 3, Dissolved Oxygen, ApparentOxygen Utilization, and Oxygen Saturation,NOAA Atlas NESDIS, vol. 63, edited byS. Levitus, 342 pp., NOAA, Silver Spring, Md.Garcia, H. E., R. A. Locarnini, T. P. Boyer, andJ. I. Antonov (2006b), World Ocean Atlas2005, vol. 4, Nutrients (Phosphate, Nitrate, Silicate),NOAA Atlas NESDIS, vol. 64, edited byS. Levitus, 396 pp., NOAA, Silver Spring, Md.Hilbrecht, H. (1996), Extant planktic foraminiferaand the physical environment in the Atlanticand Indian Oceans, Mitt. Geol. Inst. Eidg.Tech. Hochsch. Univ. Zürich, 300, 93 pp.Huber, B., J. Bijma, and K. Darling (1997),Cryptic speciation in the living foraminiferGlobigerinella siphonifera (d’Orbigny),Paleobiology, 23(1), 33–62.Jahnke, R. A. (1996), The global ocean flux ofparticulate organic carbon: Areal distributionand magnitude, Global Biogeochem. Cycles,10, 71–88.Kessler, W. S. (2006), The circulation of the easterntropical Pacific: A review, Prog. Oceanogr.,69, 181–217.Kroon, D., and K. Darling (1995), Size and upwellingcontrol of the stable isotope composition ofNeogloboquadrina dutertrei (d’Orbigny), Globigerinoidesruber (d’Orbigny) and Globigerinabulloides (d’Orbigny): Examples from the PanamaBasin and the Arabian Sea, J. ForaminiferalRes., 25, 39–53.Ku, T.-L., and T. Oba (1978), A method ofquantitative evaluation of calcite dissolutionin deep sea sediments and its application topaleoceanographic reconstruction, Quat. Res.,10, 112–129.Kucera, M., and K. Darling (2002), Cryptic speciesof planktonic foraminifera: Their effect onpalaeoceanographic reconstructions, Philos.Trans. R. Soc. London, Ser. A, 360, 695–718.Le, J., and N. J. Shackleton (1992), Carbonatedissolution fluctuations in the westernequatorial Pacific during the late Quaternary,Paleoceanography, 7, 21–42.Lea, D. W., D. K. Pak, and H. J. Spero (2000),Climatic impact of the late Quaternary equatorialPacific sea surface temperature, Science,289, 1719–1724.Locarnini, R. A., A. V. Mishonov, J. I. Antonov,T. P. Boyer, and H. E. Garcia (2006), WorldOcean Atlas 2005, vol. 1, Temperature, NOAAAtlas NESDIS, vol. 61, edited by S. Levitus,182 pp., NOAA, Silver Spring, Md.Lohmann, G. P. (1995), A model for variation inthe chemistry of planktonic foraminifera due tosecondary calcification and selective dissolution,Paleoceanography, 10, 445–457.Loubere, P. (1998), The impact of seasonality onthe benthos as reflected in the assemblages ofdeep sea foraminifera, Deep Sea Res., Part I,45, 409–432.Loubere, P. (2001), Nutrient and oceanographicchanges in the eastern equatorial Pacific fromthe last full glacial to the Present, Global Planet.Change, 29, 77–98.Loubere, P., and M. Fariduddin (1999), Quantitativeestimation of global patterns of surfaceocean biological productivity and its seasonalvariation on timescales from centuries tomillennia, Global Biogeochem. Cycles, 13,115–133.Loubere,P.,F.A.Mekik,R.François, and S. Pichat(2004), Export fluxes of calcite in the easternequatorial Pacific from the Last Glacial Maximumto present, Paleoceanography, 19,PA2018, doi:10.1029/2003PA000986.Marchitto, T. M., J. Lynch-Stieglitz, and S. Hemming(2005), Deep Pacific CaCO 3 compensation and glacial-interglacialatmospheric CO 2 , Earth Planet. Sci.Lett., 231,317–336.Mekik, F., and R. François (2006), Tracing deepsea calcite dissolution: Agreement between theGloborotalia menardii fragmentation indexand elemental ratios (Mg/Ca and Mg/Sr) inplanktonic foraminifers, Paleoceanography,21, PA4219, doi:10.1029/2006PA001296.Mekik, F. A., P. Loubere, and D. Archer (2002),Organic carbon flux and organic carbon to calciteflux ratio recorded in deep sea calcites:Demonstration and a new proxy, Global Biogeochem.Cycles, 16(3), 1052, doi:10.1029/2001GB001634.Mekik, F. A., R. François, and M. Soon (2007a),A novel approach to dissolution correctionof Mg/Ca paleothermometry in the tropicalPacific, Paleoceanography, 22, PA3217,doi:10.1029/2007PA001504.14 of 15

PA1216MEKIK AND RATERINK: CALCITE DISSOLUTION IN THE DEEP SEAPA1216Mekik, F. A., P. Loubere, and M. Richaud(2007b), Rain ratio variation in the tropicalocean: Tests with surface sediments in the easternequatorial Pacific, Deep Sea Res., Part II,54, 706–721, doi:10.1016/j.dsr2.2007.01.010.Naik, S. S., and P. D. Naidu (2007), Calcite dissolutionalong a transect in the western tropicalIndian Ocean: A multiproxy approach,Geochem. Geophys. Geosyst., 8, Q08009,doi:10.1029/2007GC001615.Ni, Y., G. L. Foster, T. Bailey, T. Elliot, D. N.Schmidt, P. Pearson, B. Haley, and C. Coath(2007), A core top assessment of proxies forthe ocean carbonate system in surface dwellingforaminifers, Paleoceanography, 22, PA3212,doi:10.1029/2006PA001337.Nürnberg, D. (1995), Magnesium in tests of Neogloboquadrinapachyderma sinistral from highnorthern and southern latitudes, J. ForaminiferalRes., 25, 350–368.Oba, T. (1969), Biostratigraphy and isotopic paleo-temperatureof some deep sea cores fromthe Indian Ocean, Second Ser. Sci. Rep. 41,pp. 129–195, Tohoku Univ., Sendai, Japan.Oppo, D., and R. Fairbanks (1989), Carbon isotopecomposition of tropical surface water during the past22,000 years, Paleoceanography, 4, 333–351.Pennington, J. T., K. L. Mahoney, V. S. Kuwahara,D. D. Kolber, R. Calienes, and F. P. Chavez(2006), Primary production in the eastern tropicalPacific: A review, Prog. Oceanogr., 69,285–317.Peterson, L. P., and W. C. Prell (1985a), Calcitedissolution in recent sediments of the easternequatorial Indian Ocean: Preservation patternsand calcite loss above the lysocline, Mar.Geol., 64, 259–290.Peterson, L. P., and W. C. Prell (1985b), Calcitepreservation and rates of climatic change: An800 kyr record from the Indian Ocean, in TheCarbon Cycle and Atmospheric CO 2 : NaturalVariations Archean to Present, Geophys. MonographSer., vol. 32, edited by E. T. Sundquistand W. S. Broecker, pp. 251–269, AGU,Washington, D. C.Richaud,M.,P.Loubere,S.Pichat,andR.François(2007), Changes in opal flux and the rain ratioduring the last 50,000 years in the equatorialPacific, Deep Sea Res., Part II, 54, 762–771,doi:10.1016/j.dsr2.2007.01.012.Rosenthal, Y., and G. P. Lohmann (2002), Accurateestimation of sea surface temperaturesusing dissolution-corrected calibrations forMg/Ca paleothermometry, Paleoceanography,17(3), 1044, doi:10.1029/2001PA000749.Rosenthal, Y., G. P. Lohmann, K. C. Lohmann,and R. M. Sherrell (2000), Incorporationand preservation of Mg in Globigerinoidessacculifer: Implications for reconstructingthe temperature and 18 O/ 16 O of seawater,Paleoceanography, 15, 135–145.Russell, A. D., B. Honisch, H. J. Spero, and D. W.Lea (2004), Effects of seawater carbonate ionconcentration and temperature on shell U, Mg,and Sr in cultures planktonic foraminifera,Geochim. Cosmochim. Acta, 68, 4347– 4361.Schiebel, R. (2002), Planktic foraminiferal sedimentationand the marine calcite budget, GlobalBiogeochem. Cycles, 16(4), 1065,doi:10.1029/2001GB001459.Schiebel, R., S. Barker, R. Lendt, and H. Thomas(2007), Planktic foraminiferal dissolution inthe twilight zone, Deep Sea Res., Part II, 54,676–686, doi:10.1016/j.dsr2.2007.01.009.Schmidt, D. N., S. Renaud, and J. Bollmann(2003), Response of planktic foraminiferalsize to late Quaternary climate change,Paleoceanography, 18(2), 1039, doi:10.1029/2002PA000831.Schmidt, D. N., S. Renaud, J. Bollmann, R. Schiebel,and H. R. Thierstein (2004), Size distributionof Holocene planktic foraminifer assemblages:Biogeography, ecology and adaptation, Mar.Micropaleontol., 50(3–4), 319–338.Spero, H., J. Bijma, D. Lea, and B. Bemis(1997), Effect of seawater carbonate concentrationon foraminiferal carbon and oxygenisotopes, Nature, 390, 497–500.Tans, P., I. Fing, and T. Takahashi (1990),Observational constraints on the global atmosphericCO 2 budget, Science, 247,1431–1438.Thompson, P. R., and T. Saito (1974), PacificPleistocene sediments: Planktonic foraminiferadissolution cycles and geochronology, Geology,2, 333–335.Toggweiler, J. R., D. Dixon, and W. Broecker(1991), The Peru upwelling and the ventilationof the South Pacific thermocline, J. Geophys.Res., 96, 20,467–20,497.Wyrtki, K. (1981), An estimate of equatorial upwellingin the Pacific, J. Phys. Oceanogr., 11,1205–1214.F. Mekik, Department of Geology, GrandValley State University, Allendale, MI 49401,USA. ( Raterink, Department of Earth andEnvironmental Sciences, Wright State University,Dayton, OH 45435, USA.15 of 15

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