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Stable isotope and trace element composition of foraminiferal calcite -

from incorporation to dissolution

Dissertation

zur Erlangung des akademischen Grades eines

Doktors der Naturwissenschaften

- Dr. rer. nat. -

im Fachbereich 2 (Biologie/Chemie)

der Universität Bremen

vorgelegt von

Bärbel Hönisch

Bremen, 2002

Tag des öffentlichen Kolloquiums:

14. November 2002

Gutachter der Dissertation:

Prof. Dieter Wolf-Gladrow

PD Dr. Ralph Schneider

Abstract

Long-term reconstruction of climate and environmental parameters from marine

sediments relies heavily on the reliability of proxy indicators from planktonic and benthic

foraminifera. Therefore, the aim of this dissertation is to improve our understanding and

confidence in planktonic foraminiferal proxies as indicators of seawater carbonate chemistry

and their stability in response to dissolution. Towards this goal laboratory experiments with

living specimens and empty shells collected from sediments were carried out.

In culture experiments with the living planktonic foraminifer Orbulina universa the

control of symbiont photosynthetic activity on the boron isotopic composition of shell calcite

was investigated (Publication I). Under low light (reduced photosynthetic rates) the boron

isotopic composition of the tests is 1.5‰ lower compared to shells grown under high light

(elevated photosynthetic rates). As boron isotope values trace seawater-pH, the lower δ11B

translates to a reduction in recorded pH of approximately 0.2 units. Data obtained for field-

grown, symbiont-barren Globigerina bulloides record significantly lighter δ11B than the

symbiont-bearing O. universa and therefore support the hypothesis that respiration and

photosynthesis are the key physiological parameters responsible for species-specific vital

effects. Although this experiment may indicate that symbiont-barren foraminifera reflect

ambient seawater chemistry more accurately than symbiont-bearing foraminifera, model

results suggest that photosynthesis- and respiration-driven offsets are constant over a wide

pH-range and do not compromise the reliability of δ11B in symbiont-bearing foraminifera as a

paleo-pH indicator.

The Ba/Ca ratio in foraminiferal shells has been shown to reflect seawater Ba

concentrations, which in turn are correlated to alkalinity. To investigate whether alkalinity

may directly influence the Ba incorporation into foraminiferal calcite and thereby compromise

the reliability of this alkalinity proxy, symbiont-bearing O. universa and symbiont-barren G.

bulloides were grown in seawater of constant Ba concentration at five different alkalinities

(Publication II). A weak negative correlation between the Ba uptake ratio in O. universa

shells and alkalinity was found under high light conditions. For an increase in alkalinity of

100 µmol kg-1 the distribution coefficient DBa (= Ba/Ca shell / Ba/Caseawater) decreased by

0.004. This change is well within the error of DBa determined to date and the weak influence

of alkalinity on Ba incorporation into foraminiferal shells is insignificant for

paleoreconstructions. Globigerina bulloides has not been calibrated for Ba before and the

experiments revealed that DBa in this species is the same as DBa in O. universa. In line with

the similar Ba/Ca uptake ratio of symbiont-bearing and symbiont-barren species, varying light

levels do not affect the Ba incorporation of O. universa.

To investigate the effect of undersaturated seawater on foraminiferal shell chemistry,

well preserved shells of the planktonic foraminifera species Globigerinoides sacculifer and

Neogloboquadrina pachyderma (sinistral coiling) were partially dissolved under controlled

conditions in the laboratory (Publication III). In addition to known dissolution effects on

Mg/Ca, δ18O and δ13C, significant effects on Sr/Ca and δ11B could be determined which are in

the same order of magnitude as observed for glacial/interglacial changes. Using previous

hypotheses to explain and discuss dissolution patterns, it becomes evident that the overall

process is not yet fully understood. While δ18O, δ13C, Mg/Ca and maybe Sr/Ca can be

explained by preferential dissolution of ontogenetic calcite and a shift of the bulk shell

chemistry to calcite secreted at greater depth (gametogenic and/or crust calcite), δ11B and

δ44Ca seem to be inconsistent with such an explanation and the dissolution patterns of these

elements need to be dominated by other processes. Reduced calcite stability due to higher

Mg/Ca was found to be insignificant to control the overall dissolution behavior in

foraminiferal shells and increasing Sr/Ca ratios demonstrate that crystal impurities are not

necessarily more prone to dissolution. The microstructural breakdown of shell surfaces, i.e.

the formation of fissures and crevices, indicates an increase in porosity upon shell corrosion

rather than the removal of outer calcite layers. The resulting increase in surface area leads to

the exposition of otherwise protected lattice areas and possibly allows certain elements to be

leached out. In summary, dissolution effects appear to be species-specific and depend on the

physico-chemical gradients encountered by vertically migrating foraminifera at different

locations.

The dissolution-driven weight loss of planktonic foraminifera shells in a defined narrow

size range has been proposed to reflect bottom water carbonate ion concentration ([CO32-]). A

number of recent studies used this relationship but disregarded a number of complications that

may limit the reconstructions. Publication IV combines experimental results on cultured

foraminifera and theoretical arguments to evaluate these complications: (1) The value chosen

for the pressure impact on the [CO32-] at calcite saturation is overestimated and should be 16

µmol kg-1 km-1 instead of 20 µmol kg-1 km-1. (2) The offset in [CO32-] between bottom and

pore water depends on the amount of organic matter that is being degraded within the

sediment and consequently the assumed constancy of the offset over time and space is highly

unlikely. (3) The initial weight of undissolved shells also changes between sites and over

geological timescales. Growth conditions such as [CO32-], light and temperature affect

respiration, symbiont photosynthesis and calcification processes and cause a significant

variability in initial shell weight. (4) As the dissolution susceptibility of foraminifera shells

varies between species, a single weight loss slope cannot be used for different species.

Correcting the published estimate of glacial bottom water [CO32-] for the various effects and

uncertainties discussed in Publication IV considerably reduces the published estimate of the

Atlantic glacial upper deep water [CO32-]-increase from +14 mol kg-1 to only +4 µmol kg-1.

Zusammenfassung

Die Rekonstruktion vergangener Klima- und Umweltbedingungen anhand von

marinen Sedimenten wird maßgeblich durch die Verläßlichkeit der Proxyindikatoren

bestimmt, die man aus Schalen planktischer und benthischer Foraminiferen gewinnen kann.

Das Ziel dieser Arbeit ist deshalb die Verbesserung des Verständnisses und Vertrauens in

Proxies aus planktischen Foraminiferen als Indikatoren mariner Karbonatchemie und der

Stabilität von Proxies in angelösten Schalen. Um mögliche Einflüsse zu untersuchen, wurden

Laborexperimente mit lebenden Foraminiferen und leeren Schalen aus Sedimenten

durchgeführt.

In Kulturexperimenten mit der lebenden planktischen Foraminifere Orbulina universa

wurde der Einfluß der Symbiontenaktivität auf die Borisotopenzusammensetzung der

Kalkschale untersucht (Publikation I). Im Vergleich zu Individuen, die unter hohem

Lichtangebot gewachsen sind (hohe Photosyntheseraten), ist die

Borisotopenzusammensetzung von Schalen, die unter geringem Lichtangebot (geringe

Photosyntheseraten) gebildet wurden, etwa 1.5‰ leichter. Da die Borisotopie ein Anzeiger

für den marinen pH-Wert ist, entspricht der geringere δ11B-Wert einer Verminderung des pH-

Werts um etwa 0,2 Einheiten. Die Borisotopenzusammensetzung der Symbionten-freien

Foraminifere Globigerina bulloides ist deutlich leichter als die der Symbionten-tragenden O.

universa und unterstützt damit die Hypothese, daß Respiration und Photosynthese die

physiologischen Schlüsselparameter für artspezifische Vitaleffekte darstellen. Obwohl dieses

Experiment nahelegen könnte, daß Symbionten-freie Foraminiferen die Meerwasserchemie

genauer aufzeichnen als Symbionten-tragende Arten, deuten Modellergebnisse daraufhin, daß

die durch Photosynthese und Respiration bedingten Unterschiede über weite pH-Bereiche

konstant sind und die Zuverlässigkeit von δ11B aus Schalen Symbionten-tragender

Foraminiferen als paläo-pH Indikator nicht beeinträchtigen.

Das Ba/Ca Verhältnis in Foraminiferenschalen spiegelt die Ba-Konzentration des

Meerwassers wider. Da die Ba-Konzentration des Meerwassers mit der Alkalität korreliert,

findet das Ba/Ca Verhältnis in Foraminiferenschalen Anwendung als Proxy für die Alkalität.

Um zu untersuchen, ob die Alkalität auch einen direkten Einfluß auf den Einbau von Ba2+ in

Foraminiferenkalk haben und damit die Zuverlässigkeit dieses Proxies beeinträchtigen

könnte, wurden die Symbionten-tragende Foraminifere O. universa und die Symbionten-freie

G. bulloides bei konstanter Ba Konzentration unter fünf unterschiedlichen Alkalitäten

gehältert. (Publikation II). Unter hohem Lichtangebot konnte dabei eine schwach negative

Korrelation zwischen Ba-Aufnahme in Schalen von O. universa und der Alkalität festgestellt

werden. Eine Zunahme der Alkalität in Höhe von 100 µmol kg-1 führt zu einer Abnahme von

0,004 im Verteilungskoeffizienten DBa (=Ba/CaForaminifere / Ba/CaMeerwasser). Dieser Unterschied

liegt innerhalb der natürlichen Varianz mit der planktische Foraminiferen Ba2+ einbauen und

ist damit unbedeutend für Paläorekonstruktionen. Über mögliche Vitaleffekte beim Einbau

von Ba/Ca in Schalen von Globigerina bulloides gab es vor dieser Studie keine Daten und die

Experimente konnten nun zeigen, daß diese Art Ba2+ im gleichen Verhältnis zur

Meerwasserkonzentration einbaut wie O. universa. In Übereinstimmung mit dem gleichen

Ba/Ca Aufnahmeverhältnis von Symbionten-tragenden und Symbionten-freien Arten zeigt

variierendes Lichtangebot keinen Einfluß auf den Ba2+-Einbau in O. universa.

Um den Einfluß von untersättigtem Meerwasser auf die Schalenchemie von

Foraminiferen zu untersuchen, wurden sehr gut erhaltene Schalen der planktischen

Foraminiferen Globigerinoides sacculifer und Neogloboquadrina pachyderma (links-

drehender Morphotyp) unter kontrollierten Bedingungen im Labor angelöst (Publikation III).

Zusätzlich zu den bereits bekannten Lösungseffekten auf Mg/Ca, δ18O und δ13C konnten

signifikante Effekte auf Sr/Ca und δ11B beobachtet werden, die in der Größenordnung von

Änderungen zwischen Warm- und Kaltzeiten liegen. Die Diskussion der gefundenen

Lösungsmuster anhand von früheren Erklärungshypothesen zeigt deutlich, daß man den

zugrundeliegenden Prozeß noch nicht vollständig versteht. δ18O, δ13C, Mg/Ca und eventuell

auch Sr/Ca können durch die bevorzugte Lösung von ontogenetischem Kalzit hinreichend gut

erklärt werden. Dabei wird das Gesamtsignal zur Chemie der äußeren Schale verschoben, die

in größeren Wassertiefen gebildet wird (gametogenetischer Kalzit und/oder Kruste). δ11B und

δ44Ca hingegen können nicht durch denselben Prozeß erklärt werden und die Lösungsmuster

dieser Elemente müssen deshalb durch andere Faktoren dominiert werden. Verstärkte

Lösungsanfälligkeit aufgrund von höheren Mg/Ca-Verhältnissen reicht nicht aus, um das

gesamte Lösungsverhalten von Foraminiferenschalen zu kontrollieren und zunehmende Sr/Ca

Verhältnisse zeigen, daß Verunreinigungen des Kristallgitters nicht notwendigerweise stärker

lösungsanfällig sind. Die Veränderung der Mikrostruktur einer Schalenoberfläche, d.h. die

Entstehung von Fissuren und kleinen Spalten, weist daraufhin, daß Lösung die Porosität

erhöht und die Schale nicht einfach schichtweise von außen nach innen abträgt. Das Resultat

ist eine Zunahme der Schalenoberfläche und damit die Exposition von Kristallgitterbereichen,

die in ungelösten Schalen vom aggressiven Meerwasser abgeschirmt sind. Damit besteht die

Möglichkeit, daß bestimmte Elemente bevorzugt herausgelöst werden können.

Zusammenfassend kann man sagen, daß Lösungseffekte artspezifisch sind und von den

physikochemischen Gradienten abhängen, die vertikal migrierende Foraminiferen an

unterschiedlichen Orten erfahren haben.

Der lösungsbedingte Gewichtsverlust planktischer Foraminiferenschalen einer definierten

Größenklasse korreliert mit der Karbonationenkonzentration ([CO32-]) des Wassers am

Meeresboden. Einige Studien haben diese Beziehung kürzlich angewendet, dabei aber eine

Reihe von Komplikationen nicht berücksichtigt, die die Aussagekraft der Rekonstruktionen

begrenzen. Publikation IV nutzt experimentelle Ergebnisse und theoretische Argumente, um

diese Komplikationen zu bestimmen und zu quantifizieren: (1) Der Wert für die

Druckabhängigkeit der [CO32-] bei Kalzitsättigung wurde zu groß gewählt und sollte 16 µmol

kg-1km-1 anstatt 20 µmol kg-1km-1 betragen. (2) Der Abbau organischen Materials im

Sediment verändert die Karbonationenkonzentration und bestimmt daher den [CO32-]-

Unterschied zwischen Boden- und Porenwasser. Räumlich und zeitlich gesehen ist es daher

höchst unwahrscheinlich, daß dieser Unterschied konstant ist. (3) Das Gewicht ungelöster

Schalen variiert zwischen verschiedenen Orten und über geologische Zeitskalen.

Wachstumsbedingungen wie [CO32-], Lichtangebot und Temperatur beeinflussen

physiologische Prozesse wie Respiration, Photosynthese der Symbionten und Kalzifizierung

und bedingen maßgebliche Unterschiede im Schalengewicht. (4) Da die Lösungsanfälligkeit

von Schalen unterschiedlicher Foraminiferenarten variiert, kann eine einzelne Beziehung

nicht universell für alle Arten angewendet werden. Korrigiert man Abschätzungen der [CO32-

]-Zunahme des glazialen Bodenwassers für die in Publikation IV diskutierten verschiedenen

Effekte, so reduziert sich die Zunahme für das obere Tiefenwasser des glazialen Atlantiks von

+14 mol kg-1 auf nur +4 µmol kg-1.

Danksagung

Das Gelingen dieser Arbeit verdanke ich der wissenschaftlichen Betreuung durch und Freundschaft mit Jelle Bijma, seinen Ideen, seinem Optimismus, Enthusiasmus und Überlebenswillen in Bor-schweren Zeiten.

Dieter Wolf-Gladrow danke ich für die Begutachtung und Betreuung der Arbeit, sowie für die Schaffung einer unvergleichlichen Arbeitsatmosphäre. Ebenso möchte ich mich bei Ralph Schneider für die Begutachtung bedanken.

I am especially grateful to Howie Spero, Ann Russell, David Lea, Dirk Nürnberg, Geert-Jan Brummer, Neven Loncaric, Abhijit Sanyal, Gary Hemming, Douglas Adams, Uli Groß, Michel Stoll and Frank Peeters. They have all contributed their time and considerable expertise to my work.

Nikolaus Gussone, Toni Eisenhauer, Silke Vetter, Folkmar Hauff und Anette Deyhle haben endlose Geduld mit immer wiederkehrenden Problemen am TIMS bewiesen und massgeblich zu den erfolgreichen Messungen beigetragen.

Similarly, this work would not have been possible without the advice and/or laboratory help and work of Pam Martin, Georges Paradis, Dotti Pak, Dave Winter and Sylvia Duncan.

I would like to thank Laurie Juranek, Megan Thomas and Heidi Iverson for their field help and a great summer on Catalina Island.

André Wischmeyer möchte ich für gute Nachbarschaft, Versorgung mit Kuchen, Schokolade und Musik danken. Ohne seine aufopfernde Unterstützung bei mathematischen und Computerfragen würde ich jetzt noch an den Problemen verzweifeln und hätte das Rennen nie gewonnen.

Richard Zeebe, Gert-Jan Reichart, Heiko Jansen und Christoph Völker danke ich für hilfreiche Diskussionen über Isotope und Spurenelemente, Lösungskinetik und Statistik.

Anja Terbrüggen gebührt besonderer Dank für die Organisation des C-Labors und seelische Aufbauarbeit bei der schlimmsten aller Laborkrankheiten: Coulometerfrust. Ebenso danke ich Frau Schwarz und Friedel Hinz für gute Zusammenarbeit.

Jan Helmke und Jürgen Pätzold danke ich für die Bereitstellung von Foraminiferen und Sedimentmaterial.

Das Überleben der Doktorarbeit bedeutet nicht nur wissenschaftliche Zusammenarbeit und Diskussion, sondern auch freundschaftliches Miteinander in der Arbeitsgruppe. Neben vorher genannten Gruppenmitgliedern möchte ich mich hier insbesondere bei Björn Rost, Albert Benthien, Ingrid Zondervan, Uta Schneider, Gerald Langer, Peter Köhler, Kai Schulz, Frank Gervais, Irini Mataliotaki, Ulf Riebesell, Uta Passow, Claudia Sprengel, Markus Geisen, Anja Engel, Markus Schartau, Ignacio Tebas, Silke Thoms und Christel Heemann bedanken.

Zuletzt, aber in mancherlei Hinsicht mehr als allen anderen, danke ich Hubertus Fischer, Uli Holzwarth und all denjenigen, die mich mit ihrer Freundschaft und ihrem Interesse in den vergangenen Jahren unterstützt haben.

Structure

This thesis is subdivided into 5 parts. Part 1 refers to the main context of this study.

Part 2 presents 4 manuscripts dealing with the main topic submitted or in preparation to be

submitted to reviewed scientific journals. To reduce repetitions, the references were excluded

from the manuscripts and combined in a separate chapter. Part 3 contains the conclusions of

this thesis and provides implications for future research. The appendix (part 4) presents the

report of working group 3 of the ESF Explanatory Workshop on "The ocean carbon cycle and

climate change", Delmenhorst, September 1-4, 2001, which deals with currently available

carbonate proxies and their major limitations.

Table of contents

1. INTRODUCTION AND MOTIVATION 1

1.1 The oceanic carbon cycle 1 1.1.1 The marine carbonate system 1 1.1.2 Carbonate chemistry in the light of biological activity 3 1.1.3 Glacial to interglacial changes in CO2 and future scenarios 5 1.2 The use of proxies in paleoceanography 6 1.3 Proxies and their limitations 11

2. PUBLICATIONS 15

2.1 Focus and outline of this study 15 2.2 Publication I: The influence of symbiont photosynthesis on the boron isotopic composition of foraminiferal shells 18 2.3 Publication II: Assessing the reliability of Ba/Ca as a tracer for alkalinity 31 2.4 Publication III: Post-depositional effects on trace metals and stable isotopes in foraminiferal calcite – Evidence from dissolution experiments 42 2.5 Publication IV: Comment to Broecker and Clark "Carbonate ion concentration in glacial-age deep waters of the Caribbean Sea" 71

3. SUMMARY AND OUTLOOK 80

3.1 Effects of symbiont photosynthesis and respiration on the stable boron isotopic composition of foraminiferal shells 80 3.2 The effect of alkalinity on planktonic foraminiferal Ba/Ca 81 3.3 Changes in planktonic foraminiferal shell chemistry after incubation in undersaturated seawater 82 3.4 Foraminifera collected from sediment cores - identifying their preservation state 83 3.5 Perspectives for future research 84

4. APPENDIX 87

Reconstructing and modeling past ocean carbonate chemistry – Working Group 3 report of the ESF Explanatory Workshop on "The ocean carbon cycle and climate change", Delmenhorst, September 1-4, 2001

5. REFERENCES 106

List of figures

page 1 Typical vertical seawater profiles of carbonate parameters....................................................4

2 Changes in surface ocean carbonate chemistry in response to increasing atmospheric

CO2.........................................................................................................................................6

3 Four species of planktonic foraminifera................................................................................8

4 The oxygen isotopic composition of marine calcites as a function of temperature and

seawater

δ18O..........................................................................................................................9

5 Boron speciation and isotope partitioning between B(OH)4- and B(OH)3 as a function of

seawater pH..........................................................................................................................10

6 Reconstructing past ocean alkalinity from foraminiferal Ba/Ca..........................................10

7 Comparison of the boron isotopic composition in shells of O. universa cultured under

HL and LL............................................................................................................................24

8 Comparison of the boron isotopic composition of the symbiont-bearing foraminifera O.

universa and the symbiont-barren G. bulloides taken from plankton tows and inorganic

carbonates.............................................................................................................................26

9 DBa in the subtropical, spinose, symbiont-bearing foraminifera Orbulina universa vs.

alkalinity...............................................................................................................................37

10 DBa in the subpolar, spinose planktonic foraminifera Globigerina bulloides compared

to alkalinity......................................... ................................................................................37

11 Comparison of Sr/Ca and Ba/Ca versus seawater pH in O. universa................................40

12 Microstructural breakdown of G. sacculifer shell surfaces monitored by scanning

electron microscopy............................................................................................................50

13 The effect of partial dissolution on Mg/Ca in G. sacculifer and N. pachyderma (sin.).......51

14 Averages of Mg/Ca in inner and outer calcite of 10 G. sacculifer shells as determined

by microprobe analysis of wall profiles...............................................................................53

15 The effect of partial dissolution on Sr/Ca in G. sacculifer and N. pachyderma (sin.).........54

16 The effect of partial dissolution on δ18O in G. sacculifer and N. pachyderma

(sin.)...........55

17 The effect of partial dissolution on δ13C in G. sacculifer and N. pachyderma (sin.)...........55

18 The effect of partial dissolution on δ44Ca in G. sacculifer and N. pachyderma (sin.).........56

19 The effect of partial dissolution on δ11B in G.

sacculifer.....................................................56

20 Schematic presentation of the life cycle of G. sacculifer: vertical migration and

varying calcification depths.................................................................................................59

21 Comparison of different dissolution rates of ontogenetic and gametogenic calcite and

the respective effect on changes of a heterogeneously distributed element in

foraminiferal calcite.............................................................................................................63

22 The effect of [CO3=] on planktonic foraminiferal shell weight............................................73

23 Foraminiferal shell weights versus pressure corrected [CO3=]...........................................77

A1 Present state of the δ11B proxy calibration..........................................................................91

A2 Increased foraminiferal shell weight under higher [CO32-] during shell growth

.................96

A3 Calcium carbonate content and calcite saturation in a modern sediment profile ...............99

List of tables

page 1 Dissolution effects on foraminiferal shell chemistry as observed in sediment

studies and laboratory dissolution experiments....................................................................13

2 Boron isotopic composition of cultured O. universa and modified seawater.......................22

3 Boron isotopic composition of plankton tow O. universa and G. bulloides........................ 22

4 Experimental Ba/Ca data for cultured shells.........................................................................36

5 Average weights of undissolved and dissolved foraminifera shells, dissolution estimates

and calcite saturation of experimental seawater....................................................................48

6 Dissolution experimental data: Minor and trace elements....................................................52

7 Numerical experiment on the dissolution susceptibility of foraminiferal Mg-calcite..........62

A1 Reconstructing past ocean carbonate chemistry: proxies, limitations and estimates..........90

A2 In situ investigation of sedimentary carbonate dissolution................................................95

General introduction 1

1. Introduction and motivation

Knowledge of the origin and amplitude of natural fluctuations in past climate systems

can be used to assess the stability of modern terrestrial and marine subsystems and their

potential range of variations in the future. Changes in the cycling of organic and inorganic

carbon in the ocean have been proposed (see for an overview Falkowski et al., 2000; Raven

and Falkowski, 1999; Sigman and Boyle, 2000) as mechanisms leading to the glacial-

interglacial changes in atmospheric carbon dioxide measured in ice-cores (Fischer et al.,

1999; Petit et al., 1999). In spite of the ocean's acknowledged importance in controlling

atmospheric carbon dioxide concentrations on glacial-interglacial timescales, the roles of

chemical and physical processes governing carbon transfers between the ocean and

atmosphere are still poorly understood. The chemical reactions determining the exchange of

CO2 between atmosphere and ocean are very complex and before we can go into the theory of

paleoceanographic reconstructions, the exchange reactions between ocean, atmosphere and

marine biosphere shall be introduced briefly.

1.1 The oceanic carbon cycle

1.1.1 The marine carbonate system The marine carbonate system encompasses the different dissolved inorganic carbon

species (CO2, H2CO3, HCO3- and CO3

2-), H+- and OH-- ions. These species are interrelated by

chemical reactions which determine their relative abundances in seawater. Following Henry's

law, gaseous CO2 dissolves into surface water directly proportional to the atmospheric partial

CO2 pressure (pCO2):

[CO2] aquatic = K0 (T, S) * pCO2, (1)

where K0 is the solubility coefficient of CO2 in seawater at a given temperature (T)

and salinity (S). The dissolved CO2 hydrates immediately with water to carbonic acid

(H2CO3), which itself dissociates to bicarbonate (HCO3-), carbonate (CO3

2-) and H+-ions:

CO2 (aq.) + H2O ⇔ H2CO3 ⇔ HCO3

- + H+ ⇔ CO32- + 2 H+ (2)

CO2 is therefore not only dissolved physically but dissociates to ionic species which

do not contribute to the aquatic partial pressure of CO2 (PCO2). This is the reason why


significantly more CO2 dissolves in seawater than any other inert gas such as nitrogen or

oxygen. Because the concentration of H2CO3 is very small, it is usually combined with CO2

(aq.) to [CO2]. For the description of the carbonate system in seawater, stoichiometric

equilibrium constants, K1 and K2, are used which are related to the ion concentrations and

depend on temperature, pressure (P) and salinity:

K1(T,S, P) =H +[ ]HCO3

−[ ]CO2[ ] (3)

K2 (T,S, P) =H +[ ]CO3

2−[ ]HCO3

−[ ] (4)

Decreasing T and S and increasing P result in a shift of the relative ion concentrations

to the left-hand side of equation (2), i.e. especially [CO32-] will decrease and [CO2] (aq.)

increase in colder, deeper and less saline waters. The sum of the dissolved inorganic carbon

species is abbreviated as ΣCO2 or DIC and defined as follows:

DIC = [CO2] + [HCO3-] + [CO3

2-] (5)

In seawater, about 90% of the DIC is present as bicarbonate, approximately 9% as

carbonate and about 1% as dissolved CO2.

Another essential quantity for the description of the carbonate system is alkalinity,

which is closely related to the electrical charge balance in the ocean. The concept of alkalinity

is anything but trivial and has been regarded and defined in many different ways (Dickson,

1981). In general, alkalinity depends on a small charge excess of conservative cations ([Na+]

+ 2[Mg2+] + 2[Ca2+] + [K+]) over anions ([Cl-] + 2[SO42-]) which is mainly compensated for

by the anions of carbonic and boric acid ([HCO3-] + 2[CO3

2-] + [B(OH)4-]). As a very good

practical approximation, total alkalinity (TA) can also be described as the sum of the charges

of the major weak acids in seawater plus the charge of OH- and minus the charge of H+.

TA ≈ [HCO3-] + 2 [CO3

2-] + [B(OH)4-] + [OH-] - [H+] ± minor constituents (6)

Analytically, total alkalinity is regarded in terms of buffer capacity, i.e. the ability to

neutralize strong acids. This property is used to quantitatively determine alkalinity by titration

with HCl.


Next to alkalinity, only DIC, pH and PCO2 can be determined analytically (for details

see DOE, 1994). As none of the carbonate system parameters varies independent from the

others, the interrelated dependency enables the oceanographer to calculate the entire

carbonate system (i.e. alkalinity, DIC, pH, PCO2, [HCO3-] and [CO3

2-]) with the knowledge

of no more than two of the constituents.

This was a very brief summary of the carbonate equilibria in the ocean. A detailed

description can be found in Zeebe and Wolf-Gladrow (2001). In the next section, we will see

how biological activity interacts with the thermodynamic equilibria just described.

1.1.2 Carbonate chemistry in the light of biological activity

One of the critical processes controlling the ocean-atmosphere CO2 exchange is

primary production in the surface ocean, and regeneration and cycling of biogenic materials in

the sea (e.g. Longhurst, 1991). Oceanic primary production takes place in the euphotic zone,

i.e. the upper layer of the ocean where sufficient light is available for photosynthesis. The

export of biogenic material from the surface to the deep ocean is called the biological carbon

pump, as it transfers inorganic carbon assimilated in the surface waters against the gradient to

the deep sea. Two biological carbon pumps can be distinguished, the organic carbon and the

inorganic calcium carbonate pump. The two pumps have opposite effects on the CO2

partitioning between ocean and atmosphere. While photosynthetically active organisms

sequester CO2 for the purpose of biomass production, the secretion of calcitic and aragonitic

skeletons by foraminifera, corals, pteropods and coccolithophores primarily increases surface

PCO2 (e.g. Frankignoulle and Canon, 1994; Wollast, 1994):

photosynthetic carbon fixation: 6 CO2 + 12 H2O → C6H12O6 + 6 O2 + 6 H2O (7)

carbonate precipitation: Ca2+ + 2 HCO3- → CaCO3 + CO2 + H2O (8)

Respiration processes in the deep ocean invert reaction (7) and release CO2 which

lowers the pH in the deep ocean and leads, in addition to the effects of higher pressure and

lower temperature, to the dissolution of calcium carbonates (reverse of reaction 8). An

example of the effects of photosynthesis, calcification, respiration and CaCO3 dissolution on

the distribution of the main dissolved constituents in seawater is displayed in Figure 1.


Figure 1: Typical vertical seawater profiles of carbonate parameters. The hydrographic station was sampled on RV PELAGIA cruise 157P in April 2000 and is located at 12°30' S and 53°68' N in the Southern Indian Ocean (Hönisch et al., unpublished data). Measured variables are DIC and alkalinity; other parameters were calculated using the CO2SYS program provided by Lewis and Wallace (1998). K1 and K2 were used as determined by Roy et al. (1993) and KSO4 as determined by Dickson (1990). In the surface layer, PCO2 (a), DIC and, to a minor extent, alkalinity (b) are reduced due to the activities of the biological community (carbon fixation and calcification processes). Carbon dioxide sequestration by photosynthesizing organisms thus leads to higher pH and [CO3

2-]. Organic matter degradation predominates below the euphotic zone, where it increases PCO2 and DIC. At 1500 m depth the aragonite saturation falls to undersaturated values (shaded area in c), i.e. aragonite starts dissolving at this depth. As a consequence, pH increases and PCO2 is lowered. The correlation between lower temperature, higher pressure and reduced carbonate saturation is also indicated in c. At this locality calcite saturation does not fall below values where calcite would dissolve within the water column.

Since calcifying organisms are often associated with symbiotic algae (foraminifera,

corals) or are themselves autotrophic (coccolithophores, coralline algae), the net effects of

photosynthesis and calcification may balance each other to a certain degree (e.g. Crawford

and Purdie, 1997; Gattuso et al., 1995; McConnaughey and Whelan, 1997; Spero and Parker,

1985). These symbiotic associations do not only affect the estimates of CO2 sinks and sources


for the ocean-atmosphere-interaction, but, as we will see later on, may also affect the

reliability of chemical recorders of past ocean conditions.

1.1.3 Glacial to interglacial changes in atmospheric CO2 and future scenarios

Observations from glacier icecores have shown that cyclic changes in atmospheric

CO2 levels occurred over the last 420,000 years with glacial periods displaying about 80

ppmv lower values compared to interglacials (~ 280 ppmv) (Fischer et al., 1999; Petit et al.,

1999). Isotope paleothermometry on the Vostok ice core revealed significant covariation

between air temperature and pCO2 of the past glacial cycles (Cuffey and Vimeux, 2001),

suggesting that CO2 may be an important forcing factor for climate. In contrast, Fischer et al.

(1999) observed that the pCO2 increase lags the warming of the last three deglaciations by

600 ± 400 years, rather arguing for an important feedback mechanism than a real climate

forcing function. However, the cyclicity between glacial and interglacial pCO2 cannot be

simply explained by higher oceanic CO2 solubility due to lower temperatures because the

concomitant sealevel decrease and salinity increase (e.g. Fairbanks, 1989) largely compensate

the pCO2 decrease due to cooling. Although many approaches have been made to determine

the major processes that control the state of the glacial ocean (e.g. Archer and Maier-Reimer,

1994; Boyle, 1988b; Broecker, 1997; Broecker and Clark, 2001b; Martin, 1990),

contradictions between theories and observations could not yet be excluded so that the

interactions between glacial-interglacial shifts in atmospheric CO2 and oceanic carbon

sequestration remain elusive (e.g. Anderson and Archer, 2002; Elderfield, 2002; Maher and

Dennis, 2001).

Understanding the origin of natural fluctuations in the past is crucial for predictions of

future variations (e.g. Stott and Kettleborough, 2002). Crowley (2000) estimated that only

about 25% of the 20th-century temperature increase can be attributed to natural variability.

Instead, most of the 20th-century warming is consistent with that predicted from green house

gas increases. Greenhouse gases absorb longwave (infra-red) radiation emitted from the earth

surface and thereby prevent the loss of solar energy to space. Concomitantly the global heat

budget increases. Atmospheric CO2 is one of those greenhouse gases. Since the industrial

revolution in the 19th-century, the atmospheric CO2 concentration has increased by >30%

from the average interglacial value of ~280 ppmv to 368 ppmv in 2000. The predictions for

the future exceed 900 ppmv by the year 2100 (Cox et al., 2000) if we do not manage to reduce

the current magnitude of CO2 emissions. The rise in atmospheric CO2 leads to changes in the

ocean carbonate chemistry (Figure 2) which could have strong impacts on the marine biota


(e.g. Gattuso et al., 1998; Riebesell et al., 2000; Wolf-Gladrow et al., 1999b), and may change

oceanic carbon uptake and cycling. The corresponding global temperature increase is

estimated to be as high as 1.4 to 5.8 K (Cox et al., 2000; IPCC, 2001). Thermal expansion and

loss of mass from glaciers and ice caps will lead to a global mean sealevel rise of 9 to 88 cm

by the year 2100 (IPCC, 2001).

7.85

7.95

8.05

8.15

8.25

5

10

15

20

25

30

35

1800 1850 1900 1950 2000 2050 2100

pH

[CO

2] and

[CO

32-] /

10 (

µmol

kg

-1)

year

pH1800 = 8.2

[CO 32-]

[CO 2]pH2100 = 7.9

Figure 2: Changes in surface ocean carbonate chemistry in response to an atmospheric CO2

increase since the beginning of the industrial revolution and future estimates according to the business as usual scenario IS92a (T = 25°C, S = 35‰). Note the different scales for [CO2] and [CO3

2-]. Figure modified after Zeebe and Wolf-Gladrow (2001).

From these numbers it is obvious that we need to quantify and understand the past to

develop a sound understanding for potential future variability. While the composition of the

paleoatmosphere can be determined quite well from ancient air bubbles trapped in polar ice

caps, no such tool exists for the ocean. To determine the physico-chemical state of the ancient

ocean, paleoceanographers rely on other measurable quantities, which will be introduced in

the following chapter.

1.2 The use of proxies in paleoceanography

In order to deduce past ocean and climate variability, the reconstruction of physico-

chemical seawater conditions is the primary objective of paleoceanography. As direct

measurements of past physico-chemical seawater conditions are no longer possible, secondary

indicators which have a close relationship to any one environmental parameter are employed


for this task. These measurable descriptors for desired (but unobservable) variables are called

"proxies" (Wefer et al., 1999).

According to Lea (1999a), proxies can be divided into three classes: biotic

components (i.e. morphologic or taxonomic climate responses such as stomata density, floral

and faunal assemblage compositions etc.), chemical tracers and physical and mineralogical

sediment properties. This study is focussed on chemical tracers which comprise the largest

proxy group. They can be organized in three sub-categories: proxies of physical seawater

properties, such as temperature; proxies of seawater composition, such as nutrient

concentration and carbonate chemistry; and proxies of sediment particle flux, such as

productivity (see also Wefer et al., 1999). Empirical relationships between proxies and their

respective environmental parameters have been established in either laboratory studies or field

calibrations. Many of these chemical proxy relationships are based on foraminifera, a group

of unicellular organisms which secrete multi-chambered calcareous shells1. Foraminifera

occur in all ocean basins and may dwell in surface waters (planktonic species) as well as on

the seafloor (benthic species). Widespread as they occur, they have the potential to record

oceanwide seawater properties. The morphologic and geometric features of their skeleton, i.e.

the arrangement of their successive chambers, enable the micropaleontologist to identify the

different species (e.g. Kemle-von Mücke and Hemleben, 1999). Each of these species favors

different environmental conditions (e.g. Bijma et al., 1990; Darling et al., 1999; Rutherford et

al., 1999) and the knowledge of these habitat preferences allows to focus paleoreconstructions

on specific locations and timescales.

This study is focussed on planktonic foraminifera (Figure 3), whose individual life

spans are on the order of 2-4 weeks (Bijma et al., 1990; Spindler et al., 1979). High

abundances in the world ocean in addition to a short reproductive cycle make foraminifera an

important contributor of biogenic calcite to open ocean marine sediments (Bé et al., 1977) and

a valuable tool for the reconstruction of past ocean conditions. Planktonic foraminifera shells

are composed of extremely pure calcite, typically about 99% by weight. The remaining 1% is

comprised of minor and trace elements such as Mg, Sr, Ba and U. Since trace elements and

different isotopes of major and minor elements are incorporated directly from seawater during

shell precipitation, shell composition reflects both seawater composition and the physical and

biological conditions encountered during precipitation. We will now see how these shell

constituents can help to elucidate past ocean carbonate chemistry.

1 Some benthic species do not actively secrete shells but collect sediment material to construct exoskeletons. For obvious reasons these species are not used as chemical recorders in paleoceanography.


living Orbulina universa living Globigerina bulloides with symbiont halo

Globigerinoides ruber Globigerinoides sacculifer

Figure 3: Four species of planktonic foraminifera; a and b are living specimens, c and d are sediment derived shells photographed by scanning electron microscopy. With the exception of G. bulloides, all shown species bear symbionts when alive. Photographies: H.J. Spero, UC Davis (a, b) and H. Hüttemann, University of Tübingen (c, d).

To define the ocean carbonate system, two elements of the system have to be known in

addition to temperature, salinity and pressure (depth). This can be any combination of pH,

alkalinity, DIC or related ion concentrations such as [CO32-]. Using stable oxygen isotopes,

the foundation of seawater temperature reconstructions was established as early as 1947 by

Urey. Following Urey's suggestion, Epstein et al. (1953) used mollusks grown in known

environments to establish a relationship linking the 18O/16O ratio of shell carbonate to the

temperature and isotopic composition of the water in which the carbonate was secreted. Since

that time numerous studies have used δ18O to deduce the oxygen isotopic composition of

seawater (e.g. Fairbanks, 1989) and calcification temperature (e.g. Emiliani, 1955). In

addition to the development of other paleotemperature proxies, recent work has also revived

interest in the use of Mg/Ca in foraminiferal calcite as a paleothermometer (Dekens et al.,

2002; Nürnberg, 1995; Nürnberg et al., 1996). Foraminiferal Mg/Ca proved especially useful

in separating the effects of temperature and salinity on δ18O (Elderfield and Ganssen, 2000;

Lea et al., 2002; Rosenthal et al., 2000). To give an example of proxy relationships, Figure 4

presents the temperature dependence of δ18O and Mg/Ca as recorded in shells of Globigerina

bulloides and in inorganic calcite.

General introduction

-2.5

-2

-1.5

-1

-0.5

0

0.5

2

3

4

5

6

7

8

14 16 18 20 22 24 26

δ18O

cal

cite

- δ18

O w

ater

(‰ V

-PD

B)

Mg/C

a (mm

ol/mol)

temperature (°C)

12-chambered shell

11-chambered shell

13-chambered shell

inorganic calcite

9

Figure 4: The oxygen isotopic composition (δ18O) of marine calcites (blue regressions, here e.g. for

the foraminifera G. bulloides: Bemis et al., 1998) is determined by a temperature dependent fractionation and the isotopic composition of seawater. Isotope values are given in δ-notation relative to a specific standard (here: Vienna Pee Dee Belemnite): δ18O = [(18O/16Osample)/18O/16Ostandard-1)*1000]. In addition to temperature and salinity effects, physiological processes may cause species-specific vital effects so that the δ18O of biogenic calcites may deviate from inorganically precipitated calcium carbonate (Kim and O'Neil, 1997). For instance, amputation of successive laboratory grown chambers of the subpolar foraminifera G. bulloides revealed an ontogenetic effect with larger shells being progressively enriched in 18O relative to smaller shells and less depleted in 18O relative to seawater (Bemis et al., 1998). Red symbols indicate Mg/Ca uptake in G. bulloides (Lea et al., 1999b). Mg/Ca is predominantly controlled by temperature, which can be used to subtract the temperature effect from foraminiferal δ18O and derive the salinity signal. Error bars are standard deviations of individual analyses of multiple amputated chambers.

With regard to specific carbonate chemistry parameters, foraminiferal Ba/Ca was used

to infer ocean alkalinity (Lea, 1993; Lea and Boyle, 1989), the stable boron isotopic

composition in foraminifera shells has become a powerful tool in reconstructing marine pH

(Figure 5, Pearson and Palmer, 1999; Sanyal et al., 1995; Spivack et al., 1993), differences

between the influence of [CO32-] on the stable carbon isotopic composition of Globigerinoides

sacculifer and G. ruber were found to be useful for past carbonate ion concentration estimates

(Bijma et al., 1999; Spero et al., 1999), and U/Ca and S/Ca have recently been suggested to

record [CO32-] as well (Erez et al., 2001; Russell, 2001; Russell et al., in prep.). Zn

concentrations in benthic foraminifera appear to correlate with bottom water carbonate

saturation (Marchitto Jr. et al., 2000). Because this study focuses especially on δ11B and

Ba/Ca, the functioning of these two proxies is briefly explained in Figures 5 and 6. General


descriptions of carbonate chemistry proxies and a discussion of their specific limitations can

also be found in the Appendix (Working Group 3 report of the ESF Explanatory Workshop on

"The ocean carbon cycle and climate change", Delmenhorst, September 1-4, 2001).

10

20

30

40

50

60

70

7 7.5 8 8.5 9 9.5 10

δ11B

(‰)

pH

seawater

modern marine carbonates

B(OH)3

B(OH)4-

b

0

100

200

300

400

7 7.5 8 8.5 9 9.5 10

conc

entr

atio

n (µ

mol

kg

-1)

pH

B(OH)4-

B(OH)3

a

Figure 5. a) Typical seawater concentrations of dissolved boron species as a function of pH (T=25°C,

S=35‰). At low pH, essentially all aqueous boron is in the B(OH)3 species while at high pH, essentially all aqueous boron is in the B(OH)4

- species. b) Because there is an isotopic fractionation between B(OH)3 and B(OH)4

-, the boron isotopic composition (δ11B) of each species is also pH dependent. The charged B(OH)4

- is supposedly the only species being incorporated in marine carbonates (Hemming and Hanson, 1992), resulting in a narrow range of δ11B in various biogenic and inorganic carbonates grown at modern seawater pH (grey box). Past changes in seawater pH are reflected in foraminiferal δ11B (Pearson and Palmer, 1999; Pearson and Palmer, 2000; Sanyal et al., 1995). Figures modified after Hemming and Hanson (1992) and Zeebe and Wolf-Gladrow (2001).

0

1

2

3

4

5

0 5 10 15 20 25 30

Ba/

Ca

in fo

ram

inife

ra s

hells

(µm

ol m

ol-1

)

Ba/Ca in seawater (µmol mol -1 )

Ba/Cashell = 0.16 * Ba/Ca seawater

a30

50

70

90

110

130

150

2280 2320 2360 2400 2440 2480

Ba

(nm

ol k

g-1

) nor

mal

ized

to S

= 3

4.7‰

alkalinity (µmol kg -1 ) normalized to S = 34.7‰

Indo-Pacific deep water

circumpolar deep water

circumpolar surface water

North Atlantic deep water

warm surface water

Ba = 0.67 * Alk -1515 R2 = 0.96

b

Figure 6. a) Planktonic foraminifera incorporate Ba/Ca proportional to the Ba/Ca ratio of seawater

(empirical relationship here for O. universa, Lea and Spero, 1992). b) Barium cycling and ocean alkalinity are similarly affected by biological uptake in surface waters and regeneration in deep waters. Although the apparent oceanwide correlation is mechanistically not well understood, changes in thermohaline circulation are supposed to redistribute Ba and alkalinity similarly, thereby allowing reconstruction of past alkalinity distributions from foraminiferal Ba/Ca (e.g. Lea, 1993).


In addition to chemical proxies of the seawater carbonate system, the preservation

state of carbonates can be used to estimate bottom water undersaturation for calcite and

aragonite. Relating the preservation state of carbonates in the sediment to the saturation state

of bottom water yields information on [CO32-]in situ. For example, the dissolution driven shell-

thinning of planktonic foraminifera (Broecker and Clark, 2001a; Lohmann, 1995) and the

CaCO3 size fraction index (Broecker and Clark, 1999) were found to approximate bottom

water [CO32-].

1.3 Proxies and their limitations The above mentioned proxies are but a small selection of the already available ones. A

remarkable range of available proxies (see, for an overview, Wefer et al., 1999) suggests that

the tools needed to do a comprehensive survey of past ocean-climate variability have already

been identified. However, many proxies bear uncertainties that complicate their interpretation.

There is abundant evidence that the incorporation of trace elements in foraminiferal calcite

does not take place according to thermodynamic equilibrium. Foraminifera, as living

organisms, actively precipitate their shells, thereby affecting both the structure and chemistry

of shell calcite. Active precipitation argues for significant biological and kinetic controls of

trace element substitution and isotope incorporation. As paleoceanographic reconstructions

can be no better than the proxies themselves, the principle task remaining is to refine and

validate these tools and ascertain which ones yield consistently reliable information.

Approaches to develop, calibrate and validate proxies are based on either field observations

(i.e. coretop sediments, sediment traps and plankton tows) or laboratory culture experiments.

Because environmental conditions often change in unison, using field data to quantify the

influence of variations in any single parameter is the more difficult approach. On the other

hand, laboratory culture experiments are limited by the lack of physico-chemical gradients

usually encountered by the vertically migrating foraminifera. However, the focus on selected

conditions is a major advantage of laboratory cultures. Culture data therefore provide an

important means by which sediment observations can be interpreted.

Beyond the uncertainties involved in specific element incorporation, one of the basic

assumptions in the use of proxies is that the primary signal remains unaltered after burial in

the geological record. However, sediment observations revealed significant variability in the

shell chemistry of planktonic foraminifera that could not be explained by oceanographic or

climatologic changes in the former habitat (Table 1). For instance, Savin and Douglas (1973)


and Bender et al. (1975) first demonstrated that planktonic foraminiferal Mg/Ca decreases

with water depth and attributed this change to partial dissolution. Subsequently, Brown and

Elderfield (1996), Rosenthal et al. (2000) and Dekens et al. (2002) have attempted to explore

how this dissolution effect varies between species and within different ocean basins.

Similarly, stable oxygen isotope compositions were demonstrated to increase in deeper

sediment cores (Erez, 1979b; Rosenthal et al., 2000; Savin and Douglas, 1973). Ignoring post-

depositional alterations, for instance with regard to Mg/Ca and δ18O, would lead to

underestimates of the real habitat temperatures. Although it seems a reasonable assumption

that enhanced porewater acidity causes the observed variability, sediment observations and

the use of dilute acid in laboratory experiments bear uncertainties which limit data

interpretation. Furthermore, many chemical proxies have not yet been regarded in the light of

selective dissolution, and quantification of the amount of foraminiferal shell corrosion - using

indices of size fraction (Berger et al., 1982; Broecker and Clark, 1999), weight loss estimates

(Lohmann, 1995; Lohmann et al., 1999), reflectance (Helmke and Bauch, 2002) and

microstructural breakdown (Bé et al., 1974) of foraminiferal shells - underlies a number of

assumptions and restrictions (e.g. Publication IV of this study). All these uncertainties limit

the quality of the proxy data base.

Much effort has already been spent on refining available proxies. However, although

remarkable progress has been made on this field, the expansion of our knowledge also raises

new questions. This study aims to contribute to the understanding of the functioning and

reliability of planktonic foraminiferal proxies.

Tab

le 1

: Dis

solu

tion

effe

cts o

n fo

ram

inife

ral s

hell

chem

istry

as o

bser

ved

in se

dim

ent s

tudi

es a

nd la

bora

tory

dis

solu

tion

expe

rimen

ts.

Ref

eren

ce

Spec

ies

Cor

e lo

catio

nδ18

Oδ13

CM

g/C

a

Sr/C

aU

/Ca

Ba/

Ca

Cd/

Ca

Wat

er∆ C

O32-

‰

/km

‰/k

m%

/km

%

/km

%/k

m%

/km

%/k

mde

pth

µmol

kg-1

SE

DIM

EN

T C

OR

ET

OP

OB

SER

VA

TIO

NS:

Pl

ankt

onic

fora

min

ifera

R

usse

ll et

al.

(199

4)G

. sac

culif

er

Cea

ra R

ise

-20

-15

3000

-470

0m39

to

-3

Rus

sell

et a

l. (1

994)

G. s

accu

lifer

O

nton

g Ja

va P

late

au0

0 ?

1600

-450

0m20

to

-15

Bro

wn

& E

lder

field

(199

6)G

. sac

culif

er

Ont

ong

Java

Pla

teau

0 0

1600

-440

0m20

to

-15

Erez

(197

9)

G. s

accu

lifer

N

orth

Atla

ntic

0.5

45

00-4

950m

-11

to-2

2 H

astin

gs e

t al.

(199

8)G

. sac

culif

er

Equa

toria

l Atla

ntic

/Car

ibbe

an0

2540

-364

5m27

to

2 Lo

rens

et a

l. (1

977)

G. s

accu

lifer

Ea

st P

acifi

c R

ise,

cen

tral P

acifi

c-1

4 -4

580-

4000

m47

to

-19

Dek

ens e

t al.

(200

2)G

. sac

culif

er

Ont

ong

Java

Pla

teau

0.2

-5

1600

-450

0m20

to

-15

Dek

ens e

t al.

(200

2)G

. sac

culif

er

Cea

ra R

ise

0.3

-7

2800

-460

0m44

to

-3

Dek

ens e

t al.

(200

2)G

. sac

culif

er

Sier

ra L

eone

Ris

e

0 ?

-14

3100

-510

0m32

to

-7

Ros

enth

al e

t al.

(200

0)G

. sac

culif

er

Ont

ong

Java

Pla

teau

0.2

-12

1600

-340

0m20

to

-3

Ros

enth

al e

t al.

(200

0)G

. sac

culif

er

Cea

ra R

ise

0.2

-6

2800

-420

0m44

to

21

Ros

enth

al e

t al.

(200

0)G

. sac

culif

er

Sier

ra L

eone

Ris

e0

-0.1

0 29

00-5

100m

32 t

o -7

Le

aet

al.(

2000

)G

.rub

erO

nton

gJa

vaPl

atea

u-1

216

00-2

500m

20to

9Er

ez (1

979)

G

. rub

er

Nor

th A

tlant

ic0.

6

4500

-495

0m-1

1 to

-22

Lore

ns e

t al.

(197

7)G

. rub

er

East

Pac

ific

Ris

e, c

entra

l Pac

ific

-8

058

0-38

00m

47 t

o -2

1 D

eken

s et a

l. (2

002)

G. r

uber

O

nton

g Ja

va P

late

au0.

2-1

4 16

00-4

500m

20 t

o -1

5 D

eken

s et a

l. (2

002)

G. r

uber

C

eara

Ris

e0.

3-5

28

00-4

600m

44 t

o -3

D

eken

s et a

l. (2

002)

G. r

uber

Si

erra

Leo

ne R

ise

0

?-7

31

00-5

100m

32 t

o -7

Lo

rens

etal

.(19

77)

N.d

uter

trei

East

Paci

ficR

ise,

cent

ralP

acifi

c-2

10

1900

-470

0m20

to-4

5D

eken

s et a

l. (2

002)

N. d

uter

trei

O

nton

g Ja

va P

late

au0.

3-2

0 16

00-4

500m

20 t

o -1

5 D

eken

s et a

l. (2

002)

N. d

uter

trei

C

eara

Ris

e0.

2-2

1 28

00-4

600m

44 t

o -3

D

eken

s et a

l. (2

002)

N. d

uter

trei

Si

erra

Leo

ne R

ise

-0.1

-16

3100

-510

0m32

to

-7

Rus

sell

etal

.(19

94)

G.t

umid

aC

eara

Ris

e-3

6-4

430

00-4

700m

39to

-3R

usse

ll et

al.

(199

4)G

. tum

ida

Ont

ong

Java

Pla

teau

-16

-20

1600

-450

0m20

to

-15

Bro

wn

& E

lder

field

(199

6)G

. tum

ida

Ont

ong

Java

Pla

teau

-10

-5.3

1600

-440

0m20

to

15

Lore

ns e

t al.

(197

7)G

. tum

ida

East

Pac

ific

Ris

e, c

entra

l Pac

ific

-16

019

00-4

700m

20 t

o -4

5 Er

ez(1

979)

G.t

runc

atul

inoi

des

Nor

thA

tlant

ic0.

945

00-4

950m

-11

to-2

2Er

ez(1

979)

G.i

nfla

taN

orth

Atla

ntic

0.6

4500

-495

0m-1

1to

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(197

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.hir

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745

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ntic

0.7

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t al.

(197

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. con

glob

atus

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st P

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c R

ise,

cen

tral P

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5 Er

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0.5

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(197

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ntic

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(197

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ntic

-0.4

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Ben

thic

fora

min

ifera

R

usse

ll et

al.

(199

4)C

. wue

llers

torf

fi C

eara

Ris

e-2

9 -1

330

00-4

700m

39 t

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M

cCor

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et a

l. (1

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C. w

uelle

rsto

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Ont

ong

Java

Pla

teau

-0.2

6

-8-1

2-2

515

00-4

500m

20 t

o -1

5

Tab

le 1

con

tinue

d: D

isso

lutio

n ef

fect

s on

fora

min

ifera

l she

ll ch

emis

try a

s obs

erve

d in

sedi

men

t stu

dies

and

labo

rato

ry d

isso

lutio

n

ex

perim

ents

. R

efer

ence

Sp

ecie

s C

ore

loca

tion

δ18O

δ13C

Mg/

Ca

Sr/C

aU

/Ca

Ba/

Ca

Cd/

Ca

Wat

er d

epth

‰‰

%%

%%

%

LA

BO

RA

TO

RY

DIS

SOL

UT

ION

EX

PER

IME

NT

S:

Plan

kton

ic fo

ram

inife

ra

H

aley

& K

linkh

amm

er (2

002)

G. s

accu

lifer

C

arib

bean

-7

1125

40m

Ben

der e

t al.

(197

5)

G. s

accu

lifer

4.

4H

önis

ch e

t al.,

this

stud

y G

. sac

culif

er

Gul

f of A

qaba

00

01.

60

00

300m

Lea

& B

o yle

(199

1)G

. tru

ncat

ulin

oide

s0

Le

a &

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le (1

991)

G. c

ongl

obat

us

Ber

mud

a R

ise

045

00m

H

aley

& K

linkh

amm

er (2

002)

O. u

nive

rsa

Car

ibbe

an-5

4-5

525

40m

H

önis

ch e

t al.,

this

stud

yN

. pac

hyde

rma

(sin

.)N

orw

egia

n Se

a0.

20.

2-2

53.

9-4

2 ?

043

?18

00m

Ben

thic

fora

min

ifera

Le

a &

Boy

le (1

993)

C

. wue

llers

torf

fi N

orw

egia

n Se

a

-1

3 ?

Lea

& B

oyle

(199

3)U

vige

rina

spp.

N

orth

wes

tern

Atla

ntic

-12

3430

m

Boy

le (1

988)

U

vige

rina

spp.

N

orth

wes

tern

Atla

ntic

2734

30m

Le

a &

Boy

le (1

993)

Ori

dors

alis

spp.

Ea

ster

n eq

uato

rial P

acifi

c

0 ?

3210

m

Cha

nges

in e

lem

ent/C

a ra

tios o

f cor

etop

sedi

men

t sam

ples

are

cal

cula

ted

as p

erce

ntag

e ch

ange

per

km

wat

er d

epth

, rel

ativ

e to

the

diss

olut

ion

onse

t in

a w

ater

dep

th p

rofil

e. S

imila

rly, i

soto

pe v

alue

s ar

e gi

ven

in a

bsol

ute

chan

ge p

er k

m. V

alue

s fo

r ∆C

O32-

(= [

CO

32-] in

situ

- [

CO

32-

] satu

ratio

n) i

ndic

ate

the

∆CO

32--r

ange

ove

r the

resp

ectiv

e de

pth

prof

ile a

nd w

ere

calc

ulat

ed u

sing

hyd

rogr

aphi

c da

ta fr

om n

earb

y W

OC

E an

d G

EOSE

CS

stat

ions

. Mar

ked

chan

ges

in s

hell

chem

istry

occ

ur w

here

∆C

O32-

falls

bel

ow 2

0 µm

ol k

g-1 (e

.g. D

eken

s et

al.,

200

2). O

bser

ved

diss

olut

ion

trend

s var

y in

tra- a

nd in

ters

peci

fical

ly.

Labo

rato

ry d

isso

lutio

n ex

perim

ents

hav

e be

en c

arrie

d ou

t und

er a

tmos

pher

ic p

ress

ure

usin

g m

odifi

ed s

eaw

ater

(Hön

isch

et a

l., th

is s

tudy

) or

dilu

te a

cid

(all

othe

r st

udie

s). W

ater

dep

th d

enot

es th

e de

pth

from

whi

ch th

e st

udy

mat

eria

l was

col

lect

ed. I

ndic

ated

val

ues

are

max

imum

ch

ange

s be

twee

n un

diss

olve

d an

d m

ost

diss

olve

d sh

ells

. N

ote

that

diff

eren

ces

in d

isso

lutio

n st

ate

exis

t be

twee

n la

bora

tory

stu

dies

. A

mbi

guou

s dis

solu

tion

trend

s are

den

oted

by

ques

tion

mar

ks.

Publications 15

2. Publications

2.1 Focus and outline of this study This dissertation reports of research on the use of planktonic foraminiferal proxies as

indicators of changes in seawater carbonate chemistry. In a first experimental series, living

planktonic foraminifera were investigated with regard to the incorporation of Ba/Ca and δ11B

as a function of physiological processes and seawater carbonate chemistry. A second set of

experiments concentrates on the preservation of foraminiferal shells and their chemical

composition after burial in the sediment.

I. The influence of symbiont photosynthesis on the boron isotopic composition of

foraminiferal shells Hönisch, B., J. Bijma, A.D. Russell, H.J. Spero, M.R. Palmer, A. Eisenhauer: The influence of symbiont photosynthesis on the boron isotopic composition of foraminiferal shells, Marine Micropaleontology, submitted 2002.

This part of the thesis investigates the reliability of δ11B as a proxy for paleo-pH and

the influence of symbiont photosynthetic activity. As microsensor studies have shown that pH

within the spine environment of planktonic foraminifera shows large variations due to

respiration and photosynthesis (Rink et al., 1998), it was investigated whether the known

boron isotopic fractionation between seawater and foraminiferal shells may be altered by

these physiological processes. The manuscript is based on laboratory experiments with living

Orbulina universa. Results of culture experiments are compared with field-grown O. universa

and Globigerina bulloides collected in plankton tows.

II. Assessing the reliability of Ba/Ca as a tracer for alkalinity Hönisch, B., A.D. Russell, J. Bijma, D.W. Lea, H.J. Spero: Assessing the reliability of Ba/Ca as a tracer for alkalinity; in preparation.

Culture experiments with the planktonic foraminifera Orbulina universa and

Globigerina bulloides have been carried out in order to investigate whether Ba- incorporation

during shell secretion is affected by seawater alkalinity. Inorganic precipitation experiments

predict such a linkage via increased in precipitation rates at higher alkalinities. As Ba2+ and

alkalinity vary proportionately in the ocean and the Ba/Ca ratio in foraminiferal shells is

assumed to reflect the seawater Ba2+ concentration, an influence of alkalinity on the Ba

incorporation could compromise the use of this proxy.

Publications 16

III. Post-depositional effects on trace metals and stable isotopes in foraminiferal calcite – Evidence from dissolution experiments

Hönisch, B., J. Bijma, N. Gussone, H.J. Spero, D. Nürnberg, D.W. Lea, A. Eisenhauer: Post-depositional effects on trace metals and stable isotopes in foraminiferal calcite – Evidence from dissolution experiments; in preparation. One of the basic assumptions in the use of proxies for paleoceanographic

reconstructions is that the primary signal remains unaltered after burial in the geological

archive. However, observations on sediment cores revealed significant variability in the shell

chemistry of planktonic foraminifera that could not be explained by oceanographic or

climatologic changes in the former habitat. Although it seems a reasonable assumption that

partial shell dissolution causes the observed variability, a number of uncertainties still remain.

Using extraordinarily well preserved shells of the tropical Globigerinoides sacculifer and the

polar Neogloboquadrina pachyderma (sin.), dissolution experiments under simulated natural

conditions have been carried out in the laboratory. Partially dissolved shells have been

analyzed with respect to minor and trace element to calcium ratios (Mg/Ca, Sr/Ca, Ba/Ca,

U/Ca, Cd/Ca) and stable isotopic compositions (δ18O, δ13C, δ44Ca, δ11B). The combination of

controlled laboratory conditions and the investigation of numerous proxies allows a detailed

discussion of the results with regard to the underlying dissolution mechanisms.

IV. The impact of the ocean carbonate chemistry on living foraminiferal shell weight: A comment to Broecker and Clark’s „Carbonate ion concentration in glacial-age deep waters of the Caribbean Sea“ Bijma, J., B. Hönisch, R.E. Zeebe: The impact of the ocean carbonate chemistry on living foraminiferal shell weight: A comment to Broecker and Clark’s „Carbonate ion concentration in glacial-age deep waters of the Caribbean Sea“. Geochemistry Geophysics Geosystems; in press.

Using the size normalized weight of planktonic foraminifera to determine their

preservation state and estimate the carbonate ion content of oceanic deep waters may be

compromised by a number of physico-chemical parameters. The assumptions made by

Broecker and Clark (2002) disregard existing evidence from culture experiments which

predict differences in shell wall thickness of planktonic foraminifera upon different growth

conditions. In addition to estimating the impact of uncertainties in the corrosivity of bottom

and pore waters, we quantify the effect of carbonate chemistry on shell growth of planktonic

foraminifera using combined data of various culture experiments.

Publications 17

Erklärung über den von mir geleisteten Anteil an den Publikationen Publikation I

Die Laborexperimente wurden von J. Bijma und mir geplant und in Zusammenarbeit mit A.

Russell und H. Spero durchgeführt. Ich habe die Proben gemessen, die Daten ausgewertet und

das Manuskript verfaßt.

Publikation II

Die Laborexperimente wurden von J. Bijma und mir geplant und in Zusammenarbeit mit D.

Lea, A. Russell und H. Spero durchgeführt. Ich habe die Daten ausgewertet und das

Manuskript verfaßt.

Publikation III

Die Laborexperimente habe ich in Zusammenarbeit mit Jelle Bijma geplant. Ich habe die

Experimente durchgeführt, ausgewertet und die Borisotope gemessen. Ich habe das

Manuskript verfaßt.

Publikation IV

Das Manuskript wurde in Zusammenarbeit mit J. Bijma und R. Zeebe geplant und verfaßt. Ich

habe die Kulturdaten zusammengestellt und ausgewertet.

Publication I 18

Publication I

The influence of symbiont photosynthesis on the boron isotopic composition of foraminiferal shells

Bärbel Hönisch, Jelle Bijma, Ann D. Russell, Howard J. Spero, Martin R. Palmer and Anton Eisenhauer

Marine Micropaleontology (submitted 2002)

.......................................................................................................................................................

Abstract Culture experiments were carried out with the planktonic foraminifer Orbulina

universa under high and low light levels in order to determine the influence of symbiont

photosynthetic activity on the boron isotopic composition of shell calcite. Under low light

(reduced photosynthetic rates) the boron isotopic composition of the tests is 1.5‰ lower

compared to shells grown under high light (elevated photosynthetic rates). In terms of inferred

pH, the lower boron isotope values correspond to a reduction in pH of approximately 0.2

units. The boron isotopic composition of Orbulina universa from plankton tows is similar to

shells grown under low light conditions in the laboratory. These data are consistent with

reduced symbiont concentrations in recently secreted shells. In addition to laboratory and

field grown O. universa, we present the first data for a symbiont-barren foraminifer,

Globigerina bulloides. Data obtained for G. bulloides fall ~1.4‰ below the field grown O.

universa. Although the plankton tow results are preliminary, they support the hypothesis that

respiration and photosynthesis are the key physiological parameters responsible for species-

specific vital effects. Model results have predicted that photosynthesis- and respiration-driven

offsets as presented here are constant over a wide pH range and thus do not reduce the

reliability of δ11B as a paleo-pH indicator.

Publication I 19

Introduction

Data from experiments with living foraminifera have confirmed the hypothesis that

seawater pH is the dominant environmental control on the 11B/10B content (δ11B) of

planktonic foraminifera shells (Hemming and Hanson, 1992; Sanyal et al., 2001; Sanyal et al.,

1996; Sanyal et al., 2000; Spivack et al., 1993). Although measurements of foraminiferal δ11B

are not yet a routine tool in paleoceanography, several studies have published paleo-pH

reconstructions across different geological timescales with encouraging results (Palmer et al.,

1998; Pearson and Palmer, 2000; Sanyal and Bijma, 1999; Sanyal et al., 1997; Sanyal et al.,

1995; Spivack et al., 1993).

Whereas pH is the primary environmental control on shell δ11B, several physiological

processes can modify the pH of the calcifying microenvironment, potentially complicating

straightforward interpretation of δ11B data. For instance, microelectrode studies have revealed

that pH in the calcifying microenvironment of symbiont-bearing foraminifera varies with light

levels (Jørgensen et al., 1985; Rink et al., 1998). Although symbionts remove CO2 during

photosynthesis, thereby increasing pH in the foraminiferal microenvironment, respiration

releases CO2 and decreases pH. Results from diffusion-reaction model simulations support

these microsensor studies (Wolf-Gladrow et al., 1999a), showing that respiration and

symbiont-photosynthesis, along with diffusion and chemical reactions, control the availability

of CO32- and HCO3

- for the calcification process. The carbonate ion effect on shell δ13C and

δ18O of planktonic foraminifera (Bijma et al., 1998; Spero et al., 1997) can also be partly

explained by the influence of these physiological processes (Zeebe, 1999; Zeebe et al., 1999).

Comparison of empirical δ11B versus pH-relationships has revealed significant offsets

between inorganic and biogenic calcification as well as between foraminifera species (Sanyal

et al., 2001). It was speculated that species-specificity could be due to differences in

microenvironment pH and/or due to differences in the relative proportion of calcite

precipitated during day and night (Sanyal et al., 2001). Similarly, Hemming et al. (1998)

attributed heavier boron isotopic compositions recorded in a coral during periods of high

productivity to enhanced symbiont photosynthetic activity and a therefore higher pH. This

study investigates the influence of symbiont photosynthetic activity on the boron isotopic

composition of O. universa grown in the laboratory. In order to estimate the effects on

naturally grown foraminifera, we compare experimental data with plankton tow samples of O.

universa and the symbiont-barren G. bulloides.

Publication I 20

2. Methods

2.1 Foraminifera collection and culturing

Foraminifera were cultured using previously established methods (Lea and Spero,

1992; Mashiotta et al., 1997; Spero et al., 1997). Juvenile (presphere) O. universa were hand

collected by scuba divers in July and August 2000 from surface waters of the San Pedro

Basin, approximately 2 km NNE of the Wrigley Institute for Environmental Studies, Santa

Catalina Island, California. Surface seawater for culturing was collected at the dive site,

filtered through a 0.8 µm membrane filter and its carbonate chemistry was subsequently

modified using the method of Sanyal et al. (2001). To reduce the number of shells required

for isotope analysis, the boron concentration in the culture solution was increased tenfold by

adding 0.27 g of boric acid (H3BO3) per L seawater. The pH was readjusted to ambient pH of

8.16 by titration with NaOH. Samples of the culture solution were taken at the beginning and

end of the experiment, acidified with ultrapure HCl and archived for later determination of the

isotopic value.

After collection, individual foraminifera were examined under an inverted light

microscope for identification of species and general condition and then transferred to 115-ml

glass jars containing the experimental filtered seawater. All culture jars were maintained at a

constant temperature in a 22 ± 0.3°C water bath, the approximate summer SST at the

collection site. For each experiment, seventy individuals were grown in the laboratory.

Foraminifera were grown under the following conditions: 1) a 12-hr high light:12-hr dark

cycle where light levels were adjusted to above Pmax (315-326 µmol photons m-2 s-1), and 2) a

12-hr low light (18-20 µmol photons m-2 s-1):12-hr dark cycle. Both experiments utilized high

output, cool white, fluorescent bulbs. The former light levels exceed the saturating irradiances

for symbionts in O. universa, whereas the latter are lower than the light compensation point

(Rink et al., 1998). During the 6-to 15-day culture period, O. universa secretes and calcifies a

spherical chamber. The foraminifera were fed a 1-day old Artemia sp. nauplius (brine shrimp)

every third day. Upon termination of the experiment following foraminiferal gametogenesis,

the empty shells were rinsed in ultrapure water and archived for later analysis.

Alkalinity was determined by Gran-titration at the start and termination of the

experiment. At the same time, dissolved inorganic carbon (DIC) samples were collected,

poisoned with saturated HgCl2 solution and measured coulometrically at the Alfred Wegener

Institute in Bremerhaven, Germany. Seawater pH values were determined potentiometrically

and are given on the NBS scale. Carbonate chemistry analyses were calibrated against

Publication I 21

certified reference material supplied by Dr. A.G. Dickson, University of California, San

Diego.

Plankton tow samples were collected at the dive site in order to determine the boron

isotopic composition of naturally grown O. universa and the symbiont-barren Globigerina

bulloides. Nets with a mesh size of 153 µm were towed at 0-20 m depth. Selected

foraminifera shells were rinsed in distilled water, dried and archived. The samples were

treated in a low temperature asher to remove organic matter and to better distinguish between

juvenile O. universa and G. bulloides. Approximately 300 shells of each species were

collected. Most O. universa had built their spherical chambers shortly before collection.

Shells were very thin and none of the collected specimens from the two species showed signs

of gametogenic calcification. Total sample weight before cleaning was no more than 1 mg for

O. universa and 0.6 mg for G. bulloides.

2.2 Analytical techniques

With the exception of the plankton tow samples, only gametogenic individuals from

the culture experiments were used for analysis. All specimens were rinsed in distilled water to

remove sea salts, dried and weighed. The shells of each experiment were pooled, crushed and

bleached with 4-6% sodium hypochlorite to remove organic matter and then rinsed,

ultrasonicated and centrifuged repeatedly with distilled water to remove soluble salt and

eventually adsorbed B. In a laminar flow bench, the cleaned carbonate was dissolved in 2N

quartz distilled HCl. The dissolved sample, containing approximately 5 ng of B, was loaded

on a rhenium zone refined filament, and 1 µl of boron-free seawater was added to enhance

ionization and suppress fractionation (Hemming and Hanson, 1994). Samples were dried at an

initial ion current of 0.8 A, followed by a 1 minute period at 1.2 A. Loaded filaments were

kept under an infrared lamp until mounted into the mass spectrometer. Isotope data were

collected on a Finnigan MAT 262 RPQ+ Thermal Ionization Mass Spectrometer (TIMS) at

GEOMAR in Kiel, Germany. The BO2- ion method was used following previously published

procedures (1997; Sanyal et al., 1996). For the culture experiments each sample was run at

least 4 times. Cultured foraminifera samples were measured at a filament temperature of 915

± 10°C. While we seldom observed time-dependent fractionation in these boron enriched

samples, the small plankton tow samples started fractionating after 20-30 minutes of

acquisition. We could therefore only complete two acceptable runs for O. universa and a

single acquisition for G. bulloides. However, initial values of the fractionating runs were

consistent with the results of acceptable analyses.

Publication I 22

Table 2. Boron isotopic composition of cultured O. universa and modified seawater

light

(µmol photons m-2 s-1)

pH

(culture water)

Seawater

δ11BMS (‰)

n

O. universa

δ11BC (‰)

n

δ11BNC

(‰)

321 ± 8 8.16 ± 0.02 -8.9 ± 0.1 5 -25.6 ± 0.6 4 22.0 ± 0.6

19 ± 2 8.15 ± 0.03 -9.1 ± 0.4 6 -27.2 ± 0.3 4 20.5 ± 0.3

Results are based on 70 shells per sample. Errors are expressed as 2σmean for multiple sample runs. δ11B (‰) = (Rs/Rstd-1)*1000, Rs = 11B/10B of sample, Rstd = 11B/10B of NBS 951 boric acid standard. Seawater standard = 39.5 ± 0.34‰. n = number of replicate analyses. δ11BNC is the δ11BC after conversion to the natural seawater scale (δ11BNS = 39.5‰), see text and Eq. 9 for details.

Table 3. Boron isotopic composition of plankton tow O. universa and G. bulloides

species ambient pH δ11B

(‰)

n

O. universa 8.12 ± 0.02 20.5 ± 0.5 2

G. bulloides 8.12 ± 0.02 19.0 ± 0.9 2*

Results are based on approximately 300 shells per sample. Errors are expressed as 2σmean for multiple sample runs. δ11B (‰) = (Rs/Rstd-1)*1000, Rs = 11B/10B of sample, Rstd = 11B/10B of NBS 951 boric acid standard. Seawater standard = 39.5±0.34‰. n = number of replicate analyses. * = runs incomplete according to criteria for acceptable runs, see text for details.

To rule out isobaric interferences on mass 42 with organic contamination (12C14N16O-

ions), mass 26 (12C14N-ions) was monitored during each measurement. No interferences were

detected. The 11B/10B ratio was corrected for isotopic interferences on mass 43 (10B16O17O-

ions) by subtraction of 0.00078 from the 43/42 ratio (Spivack and Edmond, 1986).

The fractionation ε between natural seawater (NS) and calcite (C) is usually calculated

as: ε(NS-C) = δ11BNS-δ11BC. This equation gives a good approximation when the isotopic

composition of NS and modified seawater (MS) are the same. Because the modified seawater

used in the culture experiments had a significantly different isotopic composition from natural

seawater, all analyses were corrected for this difference in order to allow comparisons to

Publication I 23

previously published data. To convert our data to the natural seawater scale we applied the

following equation (Zeebe and Wolf-Gladrow, 2001):

δ11BNC = αNS-MS * δ11BC + (αNS-MS – 1) *1000 (9)

where αNS-MS is a factor expressing the isotope difference between modified and natural

seawater. δ11BNC is the approximate value of the calcite if it had been grown in natural

seawater.

The boron isotopic compositions are listed in Table 2 and Table 3. Errors are

expressed as 2σmean. Repeated analyses of natural seawater used as a laboratory standard

resulted in an average value of 39.58 ± 0.34‰ (n = 9; filament temperature: 900 ± 10°C).

For laboratory intercomparison, additional analyses of the culture samples were

performed on a Micromass VG Sector 54 TIMS at the Southampton Oceanography Centre

(SOC), Southampton, UK. Analyses followed the method outlined in Palmer et al. (1998).

Samples and NBS 951 boric acid standard were measured at a filament temperature of 925 ±

10°C.

3. Results and Discussion

In the following section we present the data obtained from our experiments. The data

set is internally consistent and the results are reasonable with regard to theoretical

considerations. However, we found systematic offsets from previously published calibration

curves. Although the offsets do not affect the conclusions of this and most previous studies,

the underlying problem will be discussed in more detail at the end of the following section.

3.1 Laboratory experiments

The results of our experiments clearly show the influence of symbiont photosynthetic

activity on the boron isotopic composition of the shell. At equal culture water pH the δ11B of

low light O. universa shells is 1.5‰ lower than specimens grown under high light (Table 2,

Figure 7). If we shift the theoretical curve for δ11B of B(OH)4- (Kakihana et al., 1977) so it

passes through our HL data, the δ11B for the LL group would imply a decrease in pH of ~0.2

units.

Publication I 24

15

20

25

30

35

7.6 7.8 8 8.2 8.4 8.6 8.8 9 9.2

δ11B

(‰, N

BS

951)

pH (NBS scale)

theoretical curve for B(OH)4-

(Kakihana et al., 1977)

HL

LL

cultured O. universa (Sanyal et al., 1996)

2.7‰

0.2

Figure 7. Comparison of the boron isotopic composition in shells of O. universa cultured under HL (open circle) and LL (filled circle). Shells were grown in modified seawater with 10-fold increased boron concentrations. To account for the isotopic difference between culture medium and natural seawater, shell data were converted to the natural seawater scale (i.e. δ11BNS = 39.5‰, equation 9, Zeebe and Wolf-Gladrow, 2001). Also shown is the empirical HL curve for O. universa established by Sanyal et al. (1996, solid curve). Note that our data are offset to Sanyal’s values by ~ +2.7‰. In order to determine the pH at the site of calcification under LL conditions (arrow-pointed circle), we moved the theoretical curve for δ11B of B(OH)4

- versus pH (Kakihana et al., 1977), dashed line) onto our HL data point (dotted line). The reflection of the LL data at the shifted curve thus yields the pH (see arrows).

Our calculated, δ11B-based, pH offset between HL and LL conditions is smaller than the

HL-dark pH offset measured by Rink et al. (1998) using microelectrodes. Rink et al. (1998)

measured the pH within the spine microenvironment of O. universa, reporting values of 7.95

units in the dark, and 8.85 and 8.65 at 717 and 152 µmol photons m-2 s-1, respectively.

Interpolating between the latter two values we estimate a pH of 8.7 for the HL conditions in

our culture experiments (~320 µmol photons m-2 s-1). The difference between our HL

conditions and complete darkness should therefore be ~0.75 pH units. Although the pH offset

is considerably larger than our boron isotope data would predict, we note that the microsensor

data are spot measurements under specific illumination conditions. In contrast, our shell data

reflect an integrated signal over several diurnal light-dark cycles. Two studies support this

argument: (1) Spero and Parker (1985) have shown that symbiont photosynthetic rates in O.

Publication I 25

universa display a daily periodicity. For any given 12-hour illumination period, symbionts

only photosynthesize at a maximum rate for 4-6 hours with lower rates during the remaining

illuminated period. Based on symbiont density and photosynthetic rates provided in that

study, the integrated photosynthetic rate for one light period is calculated at ~59 nmol C d-1

instead of ~87 nmol C d-1 which would be calculated if the maximum photosynthetic rate had

been maintained for the full 12 hour illuminated period. Therefore, the integrated symbiont

photosynthetic effect is only 68% of the spot pH measurements made by Rink et al. (1998).

With regard to pH, the computed integrated value for a full light period is therefore only 8.46

instead of 8.70. (2) Culture experiments by Lea et al. (1995) further showed that calcification

in O. universa varies among specimens and is not strictly limited to the daylight hours. They

calculated that on average, 33% of the spherical shell is precipitated during the night. Using a

simple mass balance, the influence of combining calcite secreted during the night (@ pH =

7.95) and during the day (@ pH = 8.46) would yield a weighted, time integrated pH of 8.29

for the HL group. The pH difference predicted for foraminifera grown under a HL-dark cycle

compared to shells grown in complete darkness is therefore reduced to ~0.34 instead of ~0.75

units.

Finally, it should be kept in mind that we did not keep the LL-foraminifera in the dark

but at ~19 µmol photons m-2 s-1. Although this is below the physiological compensation point

for the O. universa symbiotic association (association respiration rate = symbiont

photosynthetic rate) (Rink et al., 1998), symbiont photosynthesis still removes CO2.

Therefore the actual microenvironment pH under LL conditions should be higher than that in

shells grown in the dark. Using this line of argument, the calculated HL-dark pH difference

of ~0.34 units should be smaller for LL grown specimens. Our experimental result of a ~0.2

pH difference between LL and HL grown specimens agrees well with these calculations.

We also compared our results to those of a diffusion-reaction model (Zeebe et al.,

2001; subm.) which predicts a δ11B offset of approximately 3‰ between an integrated diurnal

record and shells grown under complete darkness. This value is considerably larger than our

experimental result (~1.5‰) due to yet unknown reasons. Additional model runs need to be

performed to explain the difference. Irrespective of the magnitude, the most important model

prediction is the constancy of this offset over a wide seawater pH range (7.9-8.5). The use of

δ11B as a paleo-pH proxy is therefore not compromised through physiological processes in the

spine environment.

Publication I 26

16

20

24

28

32

7.6 7.8 8 8.2 8.4 8.6 8.8 9 9.2

δ11B

(‰, N

BS

951)

pH (NBS scale)

theoretical curve for B(OH)4-

(Kakihana et al., 1977)

inorganic precipitation(Sanyal et al., 2000)

O. universa

G. bulloides

Figure 8. Comparison of the boron isotopic composition of the symbiont-bearing

foraminifera O. universa and the symbiont-barren G. bulloides (filled circles) taken from plankton tows and inorganic carbonates (open circles and solid line, Sanyal et al., 2000). As the inorganic carbonate was also precipitated in artificial seawater, Sanyal et al.‘s (2000) δ11B values were converted to the natural seawater scale according to Zeebe and Wolf-Gladrow (2001, equation 9). Also shown is the reference curve for our O. universa cultured under HL conditions (dotted line, see also Figure 7 and discussion in the text) and the theoretical curve for δ11B of B(OH)4

- versus pH (dashed line, Kakihana et al., 1977).

3.2 Plankton tows

The δ11B value of O. universa collected in plankton tows (20.5 ± 0.5‰) is identical to

that of O. universa cultured under LL conditions (Table 3, Figure 8). This observation is in

contrast to sediment coretop data for this species, which were shown to be isotopically similar

to δ11B of shells grown in the laboratory at ambient pH (Sanyal et al., 1996). However,

comparison of laboratory experiments on the influence of light on the stable oxygen isotopic

composition of O. universa (Bemis et al., 1998) and plankton tow samples from the Indian

Ocean (Bouvier-Soumagnac and Duplessy, 1985) show that plankton tow samples fall

midway between the laboratory HL and LL temperature relationships. Bemis et al. (1998)

concluded that the oxygen isotopic composition observed in plankton tows is due to the fact

that the Indian Ocean O. universa were obtained from oblique tows over 70 m water depth.

Hence the tow δ18O values are due to combining shells that grew under a range of light

Publication I 27

conditions. Our plankton tow foraminifera were collected at depths down to 20 m. Light level

measurements made at the collection site in August 1987 yielded irradiance levels between

2188 µmol photons m-2 s-1 at the surface and 361 to 123 µmol photons m-2 s-1 at 18 to 27 m

water depth (Spero, unpublished data). These irradiances would suggest all the tow-collected

shells were exposed to light levels that were higher than the HL levels in the laboratory. To

explain the low isotopic value in the tows, we hypothesize that the thinly calcified specimens

collected in plankton tows are not fully calcified and may not contain the density of

symbionts expected from a similar sized sphere as it approaches gametogenesis. Thin-walled

shells could also reflect a disproportionate amount of dark (night) calcification relative to

gametogenic shells. The photosynthetic impact on the boron isotopic composition is therefore

assumed to be reduced at such an early stage suggesting plankton tow samples are not the

optimum source of O. universa material for testing the boron isotope calibration.

The δ11B of symbiont-barren G. bulloides (19.0 ± 0.9‰) was 1.4‰ lower than the O.

universa shells collected from the same plankton tows. Because this is the first δ11B datum

ever measured on a symbiont-barren species, it cannot be compared to literature data.

However, the dominant physiological process that affects the carbonate chemistry of G.

bulloides at the site of calcification is respiration. Although pH measurements have never

been conducted on this species, it is well known that the addition of respiratory CO2 decreases

pH in symbiont-bearing foraminifera by up to 0.3 units (Jørgensen et al., 1985; Rink et al.,

1998; Wolf-Gladrow et al., 1999a) and therefore would be expected to influence G. bulloides

similarly. Comparison of this datum point with data from inorganic precipitation experiments

(Sanyal et al., 2000) demonstrates that G. bulloides falls slightly below the inorganic

precipitation curve (Figure 8). Given the uncertainty of absolute differences between studies

and samples and the single datum presented here, the similarity between G. bulloides and the

inorganic precipitation experiments is promising. The lower δ11B compared to O. universa

and the inorganic precipitation results is reasonable under the assumption of a lower pH at the

site of calcification due to respiration.

3.3 Analytical offset

Our HL data are offset from Sanyal et al‘s (1996) empirical equation based on cultured

O. universa by approximately +2.7 ‰ (Figure 7). At this stage we cannot explain the offset.

Although part of the explanation could be due to lower light intensities in Sanyal’s

experiments (no special illumination was provided apart from the normal laboratory ceiling

lighting), it is unlikely that irradiances were lower than the LL levels studied in our

Publication I 28

experiments. Besides probable differences in the light regime, major differences between the

two experimental set-ups are the use of boron-enriched seawater and the fact that specimens

in our experiments underwent gametogenesis. While Sanyal et al. (2001) ruled out the

possibility that higher boron alkalinity in artificial seawater affects experimental δ11B values,

comparison between pregametogenic experimental individuals and postgametogenic shells

derived from sediments (Sanyal et al., 1996) confirmed that gametogenesis does not influence

the boron isotopic fractionation significantly. Since the experimental methods were equal

apart from these differences, there is no explanation for the offset to be expected from the

experimental point of view.

The only remaining difference is the laboratory and mass spectrometer on which the

samples were analyzed. Data for the previously published empirical relationships on

foraminifera and inorganic calcite were all established in the same laboratory (2001; Sanyal et

al., 1996; 2000). However, offsets between laboratories have already been reported in the

literature. For instance, Hemming et al. (1998) compared marine coral boron isotope data

studied by Vengosh et al. (1991), Hemming and Hanson (1992) and Gaillardet and Allègre

(1995). They found offsets up to 3‰ between studies, although measurements were

conducted on the same modern coral species. Analyses on the coral Porites (Hönisch and

Bijma, unpublished data), are similar to data published by Hemming and Hanson (1992) and

Gaillardet and Allègre (1995), indicating our analytical techniques are sound and comparable

between laboratories. Furthermore, data acquired on G. sacculifer at SOC (M.R. Palmer) are

offset by ~+2‰ to similar samples analyzed by Sanyal et al. (2001). Our own repeated

analyses on different samples of G. sacculifer revealed a much closer similarity in δ11B

between this species and O. universa than the one reported by Sanyal et al. (2001).

We suggest that the origin for the observed differences must be within the analytical

procedure. The offsets may be laboratory specific, maybe even specific for different

(biogenic) carbonates. Two possible causes of interlaboratory offsets include procedural

differences such as the temperature at which the analysis is performed, and differences in

standardization. For instance, the temperature at which the analysis is performed is species-

specific and is adjusted to the amount of boron present in the carbonate. Furthermore, many

laboratories use internal seawater standards to calibrate their data instead of the NBS 951

boric acid standard. Neither standard is a carbonate, and matrix differences may be more

important than assumed to date. The difference between the 43/42 ratio of biogenic

carbonates and seawater on the one hand, and the boric acid standard on the other, may be too

Publication I 29

large to make any of these non-carbonates a reasonable standard. There is a clear need to

define an international carbonate standard for boron isotopic analyses.

Despite the possibility of specific laboratory offsets, relative differences between

samples of the same species seem to be constant. Repeated analyses of our cultured samples

at SOC revealed a difference of ~2.2‰ between shells grown under HL (δ11B=23.9‰, n=2)

and LL (δ11B=21.7‰, n=1). Although the δ11B O. universa was measured ~2‰ heavier at

SOC compared to GEOMAR, the relative difference between the two cultured shell samples

obtained in both laboratories is the same within error. Relative comparisons with samples of

known δ11B-pH relationship are therefore feasible. However, comparison of absolute values

raised in different laboratories seems to be inappropriate until identification of the underlying

problem.

4. Conclusions

The results presented here suggest a dependence of O. universa δ11B on symbiont

photosynthetic activity similar to Hemming et al.'s (1998) observation on corals in high

productivity periods. Although the effect is significant, model results by Zeebe et al. (2001;

subm.) demonstrate that the δ11B offset is constant over a wide range of seawater pH (7.9-

8.5). Assuming that respiration and photosynthesis rates do not change with varying seawater

pH, the use of δ11B as a proxy for pH is not compromised through the vital effect presented

here.

If photosynthesis and respiration are the major parameters affecting deviations of the

shell isotopic signature from seawater pH, our results suggest that symbiont-bearing

foraminifera like O. universa and G. sacculifer should generally record higher δ11B values

and symbiont-barren species such as G. bulloides lower values compared to inorganic

calcites. Culture and field data presented here are consistent with this hypothesis but deviate

from earlier published data. In order to better understand the controls over δ11B in

foraminifera and to compare results from different laboratories, it is essential to resolve the

interlaboratory analytical offsets discussed here. Nevertheless, as long as modern samples of a

certain species are available, they can be used as a reference for ancient samples of the same

species. Using the shape of the theoretical relationship between pH and δ11B by Kakihana et

al. (1977), the differences in pH can be estimated.

Publication I 30

Acknowledgements

We thank the staff of the Catalina Marine Science Center for providing facilities to

make this work possible. We gratefully acknowledge field help by Laurie Juranek, M.

Thomas, H. Iverson and the Catalina dive crew. Invaluable was A. Sanyal’s introduction into

the secrets of δ11B analyses and boron enriched culture experiments. Of great benefit were

suggestions and comments by R. Zeebe, D. Lea, D. Wolf-Gladrow and H. Fischer and

technical discussions with A. Deyhle and N. Gussone. A review by N.G. Hemming was very

encouraging and the helpful suggestions are gratefully acknowledged. This research was

supported by DAAD grant D/00 20292 (BH), NEBROC (BH, JB), NSF grants OCE-9907044

(ADR), and OCE-9903632 (HJS). Analyses at SOC were funded by a European Commission

geochemical analytical facilities grant (BH, MRP).

Publication II 31

Publication II

Assessing the reliability of Ba/Ca as a tracer for alkalinity

Bärbel Hönisch, Ann D. Russell, Jelle Bijma, David W. Lea and Howard J. Spero

in preparation

.......................................................................................................................................................

Abstract

Here we report results of laboratory culture studies showing that seawater alkalinity

has a minor effect on the incorporation of Ba into the calcite shells of planktonic foraminifera

(Orbulina universa and Globigerina bulloides). The Ba/Ca ratio in foraminiferal shells has

been proposed as a proxy for alkalinity. If alkalinity itself has a significant influence on the

Ba incorporation into foraminiferal calcite, it would be impossible to use foraminiferal Ba/Ca

to differentiate between a change in seawater Ba2+ and a coincidental change in alkalinity.

Specimens of the symbiont-bearing species Orbulina universa and the symbiont-barren

Globigerina bulloides were grown in seawater of constant Ba2+ concentration at five different

alkalinities. The experimental alkalinity range comprised levels below, within and above the

range presently found in the ocean. We found a weak negative correlation between DBa and

alkalinity in O. universa shells under high light conditions: DBa = 0.27 (±0.04) – 4.39 (±1.58)*

10-5 * AT. For an increase in alkalinity of 100 µmol kg-1 DBa therefore decreases by 0.004.

This change is well within the error of DBa determined to date. The weak influence of

alkalinity on Ba incorporation into foraminiferal shells should not bias paleoreconstructions

based on foraminiferal Ba/Ca. Globigerina bulloides has not been calibrated for Ba before

and the experiments revealed that DBa in this species is the same as DBa in O. universa. In line

with the similar Ba/Ca uptake ratio of symbiont-bearing and symbiont-barren species, varying

light levels do not affect the Ba incorporation of O. universa.

Publication II 32

Introduction

One of the most intriguing scientific challenges today is to understand the interaction

between the atmospheric CO2 budget and the oceanic carbon cycle. Since both reservoirs are

tightly linked, information about one of them provides insight about the other. The values of

at least two carbonate system parameters are required to calculate the entire oceanic carbonate

chemistry. This can be any combination of dissolved inorganic carbon (DIC), alkalinity, pH

or related ion concentrations. Lea (1993) suggested that changes in the thermohaline

circulation should redistribute Ba2+ and alkalinity similarly, thereby allowing reconstruction

of past alkalinity distributions from benthic foraminiferal Ba/Ca. Using the modern oceanic

relationship between Ba2+ and alkalinity, Lea (1993) proposed an increase in alkalinity of

approximately 20-25 µmol kg-1 during the last glacial maximum when compared to the

Holocene. This could explain a significant amount of the glacial drop in pCO2.

Although this result is promising, recent research has shown that a number of

geochemical proxies are affected by the oceanic carbonate system. For instance, Lea et al.

(1999b) found that seawater pH influences shell Mg/Ca and Sr/Ca in the planktonic

foraminifera O. universa and G. bulloides. They attributed increased Sr/Ca under higher pH to

higher CO32- concentrations and thereby enhanced calcification rates. Similarly, shell δ13C

and δ18O of planktonic foraminifera vary inversely with the carbonate ion concentration of

seawater under conditions of constant δ13CDIC and δ18O of seawater (Bijma et al., 1998; Spero

et al., 1997). Vital as well as kinetic effects were found to explain the underlying mechanism

(Wolf-Gladrow et al., 1999a; Zeebe, 1999; Zeebe et al., 1999). More recently, Russell et al.

(2001, Russell, in prep. #548) observed a correlation between carbonate ion concentration and

U/Ca in foraminifera shells. If Ba incorporation into shell calcite is similarly affected, then it

would negate the use of this proxy for estimating alkalinity in paleoceans.

The specific motivation for this investigation is based on similarities in Ba2+ and Sr2+

incorporation into inorganic and biogenic calciumcarbonates. Ba and Sr are metals with ionic

radii greater than Ca and Mg (Ba2+: 1.47 Å, Sr2+: 1.31 Å, Ca2+: 1.00 Å, Mg2+: 0.72 Å,

according to Shannon (1976). Like Mg2+, Ba2+ and Sr2+ substitute for Ca2+ in biogenic calcite

(Lea and Boyle, 1991; Mackenzie et al., 1983; Speer, 1983). A general tendency has been

observed according to which partition coefficients of metals with ionic radii smaller than Ca

decrease with increasing precipitation rate, whereas those with an ionic radius larger than

Ca2+ increase with precipitation rate (Lorens, 1981). Inorganic precipitation experiments have

shown that the partition coefficients of Sr2+ and also Ba2+ depend strongly on precipitation

rate (Lorens, 1981; Morse and Bender, 1990; Tesoriero and Pankow, 1996) and observations

Publication II 33

on Sr2+ in planktonic foraminifera (Lea et al., 1999b), coccolithophorids (Stoll et al., 2001)

and apparently also in benthic foraminifera (Elderfield et al., 1996) are in agreement with

these inorganic experiments. Especially in the case of planktonic foraminifera Lea et al.

(1999b) showed that Sr/Ca increases in shells grown under higher pH. Under inorganic

conditions the adsorption of Ba and Sr (when both equally concentrated in the aqueous

solution) onto solid calcite is low but comparable (Zachara et al., 1991). Despite large

differences in seawater ion concentration, the relative uptake of Sr and Ba into planktonic

foraminiferal calcite (as reflected in the values of DSr and DBa) is similar (Lea, 1999b).

These similarities in the geochemical behavior of Ba and Sr led us to question whether

the incorporation of Ba2+ into biogenic calcite is affected by seawater carbonate chemistry.

Using laboratory experiments, we explore the influence of alkalinity on Ba/Ca in living

planktonic foraminifera.

Experimental Methods Collection and culturing of foraminifera

Foraminifera were cultured using previously established methods (Lea and Spero,

1992; Mashiotta et al., 1997; Spero et al., 1997). Juvenile (presphere) Orbulina universa and

small Globigerina bulloides were hand collected by scuba divers in July and August 2000

from surface waters of the San Pedro Basin, approximately 2 km NNE of the Wrigley

Institute for Environmental Studies, Santa Catalina Island, California. Surface seawater for

culturing was collected at the foraminifera collection site and filtered in the laboratory using

acid-cleaned 0.4 µm polycarbonate membrane filters and an acid-leached polysulfone filter

holder. After collection, individual foraminifera were examined under an inverted light

microscope for identification and inspection of general condition and then transferred to 120

ml glass jars containing the filtered culturing solutions. To avoid contamination of culture

water during transfer and feeding of specimens, sample handling was done wearing powder

free gloves and using acid-leached glass pipettes for feeding and transferring foraminifera. All

culture jars were maintained at a constant temperature in a 22 ± 0.1°C water bath, the

approximate summer SST at the collection site.

Foraminifera were grown under the following conditions: 1) a 12-hr high light

(HL):12-hr dark cycle where light levels were adjusted to above Pmax (299-365 µmol photons

m-2 s-1), and 2) a 12-hr low light (LL):12hr-dark cycle (23 µmol photons m-2 s-1). These light

The Geology Department

In the paragraph comparing the geochemistry of Ba and Sr you don’t specifically mention the influence of pH or CO3 on Sr/Ca incorporation \(aside from the precipitation rate issue\). I’d stick that in \(see Lea, Mashiotta, and Spero, 1999\)

Publication II 34

levels either exceed the saturating irradiances for symbionts in O. universa or fall below the

compensation point for photosynthetic activity (Rink et al., 1998).

During a 7- to 10-day culture period O. universa secretes and calcifies a spherical

chamber, whereas G. bulloides forms between two and four new chambers. Globigerina

bulloides and O. universa were fed a 1-day old Artemia sp. nauplius (brine shrimp) every

other or every third day respectively. After the foraminifera underwent gametogenesis, empty

shells were rinsed in ultrapure water and archived for later analysis. In addition, a sample of

the culture solution was acidified and archived to verify that the Ba2+ concentration remained

constant over the course of the experiment. Because the amount of calcite precipitated by

foraminifera does not require more than 0.1% of the initial Ca2+ present in the culture

seawater, Ca2+ concentrations were constant over the duration of the experiments.

Total alkalinity (AT) was modified by the addition of ultrapure HCl to lower AT or the

addition of ultrapure NaOH to increase AT. Initial and final alkalinity was determined by

Gran-titration with a Metrohm 785 titrino auto-titrator. Samples for DIC analysis were taken

at the beginning and the end of the experiment, poisoned with a few drops from a saturated

HgCl2 solution and measured coulometrically at the Alfred Wegener Institute.

Sample preparation

Only gametogenic individuals were used for analysis. All specimens were rinsed in

distilled water to remove sea salts, dried and weighed. Chambers secreted in the laboratory,

i.e. under controlled conditions, were separated from the field grown shell. Separating and

cleaning procedures followed methods established by Mashiotta et al. (1997). Briefly, shells

of O. universa were cracked open with a disposable scalpel and the juvenile test (if present)

was removed with a small brush. The fragments were then transferred to 0.5 ml

polypropylene centrifuge vials. Individual shells of O. universa shells were analysed where

possible, but for the smallest individuals, two or three shells were pooled to obtain at least 40

µg of uncleaned calcite. For G. bulloides, chambers grown in the laboratory were identified

by comparing the size of the specimen at collection with the size of the postgametogenic

shell. The laboratory-grown chambers were removed with a scalpel, pooled (25-35 chambers

per sample) and loaded into 0.5-ml polypropylene centrifuge vials. Samples were then subject

to a series of physical and chemical treatments including: oxidation in hot H2O2 – NaOH

solution to remove organic matter, 2-3 weak acid leaches (0.001N HNO3) and repeated rinses

in ultrapure water.

Publication II 35

Analytical Methodology

Sample analysis followed the multi-element inductively coupled plasma mass

spectrometry (ICP-MS) method previously described by Lea and Martin (1996). After

cleaning, 20-30 µg of purified foraminifera shells were dissolved in 0.5 ml of a 0.1 N HNO3

solution containing calibrated concentrations of 135Ba and 45Sc. The solutions were aspirated

into a Finnigan Element 2 high resolution magnetic sector ICP-MS. The 135Ba/138Ba and 45Sc/48Ca ratios are determined by pulse and analog acquisition modes, respectively. The

concentrations of Ba2+ and Ca2+ are then determined by isotope dilution and internal standard

calculation, respectively. Na/Ca was determined to be certain that the hydrogen peroxide –

sodium hydroxide solution used in the sample preparation was completely rinsed out. Several

analyses of a consistency standard with Ba2+ and Ca2+ concentrations similar to the

foraminiferal samples had a standard deviation of 0.4% for Ba2+, 1% for Ca2+, and 1.25% for

the Ba/Ca ratio. Analyses were performed at the University of California, Santa Barbara.

An average of 2-4 replicates was determined on O. universa. Due to the small sample

yield in G. bulloides merely one Ba/Ca analysis per alkalinity could be realised for this

species.

Results

Fifteen water samples were randomly selected from 184 total culture experiments to

verify the constancy of the trace and minor element concentrations over the course of the

experiment. These samples yielded ambient seawater concentrations of 37.8 ± 0.35 nmol kg-1

and 10 ± 0.09 mmol kg-1, for Ba2+ and Ca2+ respectively. The amount of Ba2+ incorporated

into the foraminiferal shell is negligible compared to the total Ba2+ present in the culturing

solutions (Lea and Spero, 1992). However, problems could arise from barium contamination

(e.g. during feeding) or adsorption onto the container walls. Although one of the water

samples showed a substantially higher Ba2+ concentration by 11%, the average change in the

seawater Ba2+ concentration was only 0.9% indicating that barium adsorption was negligible

and contamination was unlikely. Nevertheless, we cannot rule out the possibility that some of

the foraminifera experienced Ba2+ concentrations that may have differed slightly from the

average of 37.8 nmol kg-1.

Publication II 36

Table 4. Experimental data for cultured shells

experiment

alkalinity µmol kg-1

light level

µmol photons m-2s-1

Ba/Ca

µmol mol-1

sample weight

µg

n

DBa

Orbulina universa

BH-3 2047 332-349 0.76 52 3 0.189 ± 0.017 0.67 50 2 0.60 65 1

BH-6 2122 299 0.57 46 2 0.165 ± 0.018 0.70 59 2 0.61 39 1

BH-1 2253 23 0.70 58 3 0.164 ± 0.015 0.57 37 2 0.59 40 2 0.63 48 1

BH-2 2268 365 0.71 51 2 0.177 ± 0.01 0.67 54 2 1.06 * 45 1 0.65 114 1

BH-5 2436 332-365 0.64 56 1 0.163 ± 0.014 0.56 40 1 1.01 * 40 1 0.56 68 1

BH-4 2632 299 0.58 69 2 0.154 ± 0.004 0.60 39 3 0.57 57 1

Globigerina bulloides

BH-3 2047 332-349 0.69 6.2 25 0.183 BH-1 2253 23 0.65 34.1 35 0.173 BH-2 2268 365 0.55 14.7 25 0.146 BH-4 2632 299 0.70 26.6 25 0.184

n = number of pooled shells per analysis; * = rejected

Ba/Ca-variability within an experimental group is larger than the analytical error in

measuring the ratio. Replicate samples were excluded from means and figures (but not from

tables) if the Ba/Ca differed by more than 2σ from the experimental mean Ba/Ca. Using the

experimental seawater concentrations of Ba2+ and Ca2+ to calculate the partition coefficient

for O. universa yields a linear decrease of DBa from 0.189 ± 0.017 at AT=2047 µmol kg-1 to

0.154 ± 0.004 at AT=2632 µmol kg-1 (Fig. 9, Table 4). Statistical analysis using an ANOVA

variance ratio test indicates the DBa decrease with increasing alkalinity is significant at the

90% confidence level. Linear regression yields:

DBa = 0.27 (±0.037) - 4.39 (±1.58) * 10-5 * AT R2 = 0.37

Publication II 37

0.12

0.14

0.16

0.18

0.20

0.22

2000 2100 2200 2300 2400 2500 2600 2700

O. universa

DB

a

alkalinity (µmol kg -1 )

DBa

= 0.27 (±0.037) - 4.39 *10-5 (±1.58 *10 -5)* AT

R2 = 0.37

Figure 9. DBa in the subtropical, spinose, symbiont-bearing foraminifera Orbulina universa vs. alkalinity. Closed symbol indicates low light conditions. Each point is the average of two to four analyses of either individual or 2-3 pooled spheres (Table 4). Error bars are standard deviations of replicate analyses. The regression is a linear fit through the individual O. universa points at saturating light levels.

0.12

0.14

0.16

0.18

0.20

0.22

2000 2100 2200 2300 2400 2500 2600 2700

G. bulloides

DB

a

alkalinity (µmol kg -1 )

Figure 10. DBa in the subpolar, spinose planktonic foraminifera Globigerina bulloides compared to alkalinity. Each point is an individual analysis of multiple amputated chambers (Table 4).

Publication II 38

DBa in G. bulloides ranges between 0.146 and 0.184 (Fig. 10). These partition

coefficients are comparable to the range observed in O. universa shells over the same

alkalinity range. However, given the small number of data points and non-linear distribution

of these data, it is not possible to determine whether or not a relationship between alkalinity

and DBa exists in G. bulloides.

In a final experimental series, we investigated the influence of symbiont

photosynthetic rate (as it varied with light levels) on O. universa DBa in ambient seawater (AT

= 2257 µmol kg-1). Under LL conditions DBa= 0.164 ± 0.015 was slightly lower than under

HL where DBa= 0.177 ± 0.01 (Fig. 9, Table 4). However, these partition coefficients are not

significantly different from each other.

Discussion

Recordability of Ba/Ca in planktonic foraminiferal shells

The average partition coefficients for Ba/Ca in O. universa and G. bulloides found in

this study are DBa= 0.166 ± 0.01 and DBa= 0.171 ± 0.02, respectively. Lea and Spero (1994)

calculated a value of DBa= 0.16 ± 0.01 (regression forced through zero) for cultured O.

universa. The combined partition coefficients of core-top fossil planktonic foraminifera yield

DBa= 0.19 ± 0.05 (Lea and Boyle, 1991). Hence, our values are in good agreement with the

literature.

Globigerina bulloides has not been calibrated for Ba before. Although this species is

symbiont-barren, the DBa is similar to that of the symbiont-bearing O. universa and G.

sacculifer. This finding is in contrast to other elements which are incorporated to different

degrees by symbiont-bearing and symbiont-barren species. For instance, Mashiotta et al.

(1997) observed strong contrasts between O. universa and G. bulloides in the uptake of Cd2+.

They suggested that symbiotic dinoflagellates either influence foraminiferal incorporation of

Cd2+ by sequestering Cd2+ from the calcifying microenvironment of O. universa, or by

photosynthetically enhancing calcification rate leading to Cd2+ exclusion. In contrast, Ba has

no reported biochemical function and in line with the similarity of HL and LL results for O.

universa, the uniformity of our results suggests that symbionts have no influence on the Ba

uptake.

In order to examine DBa in relation to calcification rate we refer to Mashiotta et al.

(1997) who estimated higher calcification rates for O. universa (~3000 µmol m-2 h-1, Lea et

Publication II 39

al., 1995) than for G. bulloides under identical culture conditions (~1700-2400 µmol m-2 h-1).

This difference is not reflected in the uptake of Ba during shell formation and the similarity in

DBa between O. universa and G. bulloides supports the generally held assumption that

species-specific differences are low for planktonic foraminifera and vital effects are not

important for the Ba uptake.

The observed negative correlation between Ba/Ca and alkalinity was not expected

from theoretical arguments which would have predicted a kinetic effect towards higher Ba/Ca

with increasing precipitation rate. To better compare the magnitude of the observed effect

with results on Sr/Ca, we also plotted Ba/Ca versus pH. Total CO2 was constant in our

experiments so that pH and [CO32-] varied almost linearly with the experimental alkalinity

range. While the kinetic effect documented for the pH-dependency of Sr/Ca is ~+1% per 0.1

pH unit (Fig. 4, Lea et al., 1999b), Ba/Ca decreases by ~2% per 0.1 pH unit (Fig. 11b). The

observed effect is thus similar in magnitude but opposite in direction.

The geochemical similarities between Sr and Ba discussed earlier do not extend to the

influence of alkalinity (or pH). This may be due to dissimilarities in sorptive behavior because

of differences in aqueous complexation (e.g. (van Cappellen et al., 1993). Alternatively, since

the mineral growth in foraminifera appears to follow a pattern which is predetermined by the

molecular organisation of the organic template (Langer, 1992), it may be that the organic

matrix controls the selection of certain divalent cations and discriminates against others.

Unfortunately the process of chamber formation and calcification in foraminifera (e.g. Bé et

al., 1979; Hemleben et al., 1977; Spero, 1988) is not very well understood at this level so that

this consideration remains rather speculative.

Paleoceanographical significance

The negative slope determined for O. universa under HL conditions implies that the

error yields a change in DBa of 0.004 (~ 0.017 µmol mol-1 in Ba/Ca) if alkalinity were to

change by 100 µmol kg-1. With such a large alkalinity change, ∆DBa would still be well within

the analytical error of the partition coefficient. For comparison, Lea (1993) found an increase

of 1.8 µmol mol-1 (~50 %) in the Ba/Ca of Circumpolar Deep Water during the last glacial

period. The alkalinity-dependency of Ba/Ca (<1 % per 10 µmol kg-1) would be negligible for

the proposed changes in alkalinity on glacial/interglacial time scales. Even under glacial

Publication II

1.10

1.15

1.20

1.25

1.30

1.35

1.40

1.45

1.50

7.5 8 8.5 9

Sr uptake vs. seawater pH

Sr/C

a (m

mol

/mol

)

pH (NBS)

Sr/Ca = 0.18 + 0.138 * pH R2 = 0.46

a

40

0.45

0.50

0.55

0.60

0.65

0.70

0.75

0.80

0.85

7.5 8 8.5 9

Ba uptake vs. seawater pH

Ba/

Ca

(µm

ol/m

ol)

pH (NBS)

Ba/Ca = 1.56 - 0.113 * pH

R2 = 0.38

b

Figure 11. a) Sr/Ca in O. universa. Data were taken from Lea et al. (1999b). Shell Sr/Ca increase approximately 1% per 0.1 pH units. b) Ba/Ca in O. universa versus seawater pH. Shell Ba/Ca decrease approximately 2% per 0.1 pH units. Theoretical arguments would have predicted an increase similar to the one observed in Sr/Ca.

conditions, when pH in surface waters may have increased by 0.2 pH units (Sanyal et al.,

1995), alkalinity would not be expected to exceed the range investigated in this study. In

contrast, the deep water value of +0.3 pH units as determined by Sanyal et al. (1995) for the

LGM would require an increase in alkalinity by more than 300 µmol kg-1 in order to bring

estimated pH and observed atmospheric pCO2 values into line. Alkalinity changes on this

Publication II 41

order could lead to a potentially large Ba/Ca change. However, the deep water pH estimate is

based on a sample of mixed benthic foraminifera species and as such not very reliable. In fact,

recent data by Anderson and Archer (2002) argue against such a strong increase in pH.

Conclusions

Our results demonstrate that species variability and symbiont effects are not important

for planktonic foraminifera. Orbulina universa and G. bulloides incorporate Ba in a similar

ratio to seawater concentration. Water column alkalinity and Ba/Ca are at most weakly linked

for these species, and the effect is too small to be significant for paleoceanographic

reconstructions. Alkalinity, the closest related component to the oceanic barium cycle, can

therefore be extensively ignored as a potential factor directly affecting the incorporation of

Ba2+ into planktonic foraminiferal calcite.

Acknowledgements

We thank the staff of the Catalina Marine Science Center for providing facilities to

make this work possible. We gratefully acknowledge field help by Laurie Juranek, Megan

Thomas, Heidi Iverson and the Catalina dive crew. The experiments could not have been

successfully completed without the laboratory work of Georges Paradis and Dotti Pak.

Support and comments from Christoph Völker and Hubertus Fischer are greatly appreciated.

This research was supported by DAAD grant D/00 20292 (BH), NEBROC (BH), NSF grant

OCE-9907044 (ADR), OCE-9906821 (DWL), OCE-9729203 and OCE-9903632 (HJS).

Publication III 42

Publication III

Post-depositional effects on trace metals and stable isotopes in foraminiferal calcite – Evidence from dissolution experiments

Bärbel Hönisch, Jelle Bijma, Nikolaus Gussone, Howard J. Spero, Dirk Nürnberg, David W. Lea and A. Eisenhauer

in preparation

.......................................................................................................................................................

Abstract

Sediment observations and laboratory dissolution experiments have increased

paleoceanographers' interest in post-depositional changes of chemical proxies recorded in

foraminifera shells. While previous studies are limited by a number of uncertainties, young

proxies like δ11B and δ44Ca have not yet been regarded in the light of dissolution. We used well

preserved planktonic foraminifera species Globigerinoides sacculifer and Neogloboquadrina

pachyderma (sinistral coiling) to partially dissolve them under controlled conditions in the

laboratory, simulating the natural processes at the seafloor. In addition to known dissolution

effects on Mg/Ca, δ18O and δ13C, we observe significant effects on Sr/Ca and δ11B on the order

of glacial/interglacial changes. Discussing previously hypothesized explanations for partial

dissolution, it will be shown that the overall process is not yet fully understood. While δ18O,

δ13C, Mg/Ca and maybe Sr/Ca can be explained by preferential dissolution of ontogenetic calcite

and a shift of the bulk shell chemistry to calcite secreted at greater depth, δ11B and δ44Ca seem to

be inconsistent with such a process and other explanations have to be investigated instead. Mg

impurities were found to be insufficient to reduce the overall stability of certain foraminiferal

shell components and increasing Sr/Ca ratios demonstrate that crystal impurities are not

necessarily more prone to dissolution. The microstructural breakdown of shell surfaces is

systematic and may be useful to estimate the preservation state of shells deposited in sediments.

Dissolution drives deep fissures into the shell wall, indicating that corrosion does not only

remove outer layers of a shell but also increases shell porosity. The resulting exposition of

Publication III 43

otherwise protected lattice areas possibly allows certain elements to be leached out. Dissolution

effects appear to be species-specific and depend on the physico-chemical gradients encountered

by vertically migrating foraminifera at different locations.

Introduction

The chemical composition of biogenic carbonates derived from marine sediments is

routinely used for paleoceanographic and paleoclimatic reconstructions. One of the basic

assumptions is that the primary signal remains unaltered after burial in the geological archive.

However, observations on sediment cores reveal significant variability in the shell chemistry of

planktonic foraminifera that can not be linked to oceanographic or climatologic changes in the

organism’s habitat. Part of the observed variability is assumed to be caused by dissolution. In

order to test this assumption, dissolution studies have focussed primarily on two experimental

approaches: (1) empirical observations on carbonate samples deposited in sediments at various

depths (e.g. Berger, 1967; Brown and Elderfield, 1996; Honjo and Erez, 1978; McCorkle et al.,

1995; Zachos et al., 1994); and (2) the determination of CaCO3 dissolution kinetics and changing

chemical composition in the laboratory (e.g. Bé et al., 1974; Bender et al., 1975; Boyle, 1988a;

Caron et al., 1990; Lea and Boyle, 1991; Lorens et al., 1977). Based on these and other

experiments, several hypotheses have been proposed:

1. Shell Mg increases the susceptibility to dissolution (Brown and Elderfield, 1996; Lorens et al.,

1977; Savin and Douglas, 1973).

2. Dissolution selectively removes the more soluble individuals and thus reduces their

contribution to the chemical bulk signal (Erez, 1979b).

3. Elements that are distributed homogeneously throughout a shell are not affected by dissolution

(Bender et al., 1975; Lea and Boyle, 1991; McCorkle et al., 1995).

4. Selective dissolution of ontogenetic calcite shifts bulk Mg/Ca and δ18O toward the chemistry

of the secondary crust (Brown and Elderfield, 1996; Dekens et al., 2002; Lohmann, 1995;

Rosenthal et al., 2000; Russell et al., 1994).

5. Although dissolution alters Mg/Ca and δ18O, the initial relationship remains constant such

that paired analyses of δ18O and Mg/Ca can be used to reconstruct the δ18O of seawater

(δ18Ow) (Rosenthal et al., 2000).

Publication III 44

Although it seems a reasonable assumption that dissolution causes the observed

variability, the impact of dissolution has not yet been quantified. Whereas interpretation of

results from sediment studies contains uncertainties about the real dissolution state,

recrystallization processes and mixing of shells from different time intervals, the use of dilute

acid in laboratory dissolution experiments results in a much more aggressive dissolution than at

the seafloor. We have carried out dissolution experiments under controlled laboratory conditions

in natural seawater whose carbonate chemistry has been manipulated. Demonstrably well

preserved shells of the two planktonic species Globigerinoides sacculifer and Neogloboquadrina

pachyderma (sinistral) have been partially dissolved. From these shells we present data on

Mg/Ca and Sr/Ca as well as stable isotope values such as δ18O, δ13C, δ11B and δ44Ca. In addition

to the bulk shell composition determined by ICP-MS (inductively coupled plasma mass

spectrometry), an electron microprobe was used to measure Mg/Ca in different parts of the

calcite of individual tests.

Methodology

Samples

We selected two sediment cores with unusually well preserved foraminifera shells: The

site of sediment core M 23352 (70°0‘N, 12°25‘W, 1819 m) is thought to have been situated

above the lysocline during the past 5 glacial-interglacial cycles. The small sieve size fraction

(279-309 µm) of the polar planktonic species N. pachyderma (sin.) was hand-picked from the 57

cm sediment horizon, which corresponds to marine isotope stage (MIS) 2. Scanning electron

microscopy (SEM) examination, as well as weight and reflectance analyses revealed that these

glacial foraminifera shells are perfectly preserved (Helmke and Bauch, 2002).

Box core GeoB 5815 (29°30,64‘N, 34°58,98‘E) was collected from a water depth of 326

m. At this locality, water column and pore waters are strongly supersaturated with respect to

calcite as reflected in the abundant occurrence of aragonitic pteropod shells within the sediment.

Because the sedimentation rate in this region is 20-30 cm ka-1 [pers. comm. J. Pätzold,

University of Bremen], one can infer that the samples are mostly Holocene and contamination

with glacial individuals is rare. Globigerinoides sacculifer was hand-picked from the 10-20 cm

sediment horizon. The shells were taken from the 315-450 µm and 450-500 µm sieve size

fraction. To remove clay and other adhering particles, the shells were cleaned ultrasonically in

artificial seawater (pH 8) and subsequently rinsed in ultrapure water (pH >8) to avoid the

crystallization of salts.

Publication III 45

Experimental Methods

The dissolution experiment presented here was designed to simulate natural conditions at

the seafloor. Adjusting the saturation state under laboratory conditions (15°C, atmospheric

pressure) was accomplished by reducing alkalinity and pH at natural concentrations of dissolved

inorganic carbon (DIC): Ultrapure 9.5 N HCl was added to 0.45 µm filtered North Sea water at

2137 µmol kg-1 DIC, and 2348 µmol kg-1 alkalinity. Four saturation states were prepared: 120,

90, 70 and 50% with respect to calcite (Table 5). Carbonate chemistry was monitored by

measuring DIC and alkalinity for the initial water batches and finally for each bottle separately.

Measurements were done coulometrically and by Gran titration, respectively. Saturation state

calculations are based on the CO2SYS program provided by Lewis and Wallace (1998), using K1

and K2 as determined by Roy et al. (1993) and KSO4 as determined by Dickson (1990). The two

size fractions of G. sacculifer and the shells of N. pachyderma (sin.) were incubated in individual

experiments. Due to dissolution in the closed system, alkalinity and calcite saturation increased

during the experiment.

Whole specimens were selected from the ultrasonically cleaned foraminifera shells for

the dissolution experiments. For G. sacculifer, five subsamples of similar weight and a fixed

number of shells (i.e. 35 shells of the large fraction and 100 shells of the small fraction) were

prepared for each saturation state. For the determination of the boron isotopic composition two

larger samples with 500 shells each were selected from the small G. sacculifer fraction and

incubated at 70 and 50% saturation. For N. pachyderma (sin.), no more than 600 shells were

available. These were split into groups of 100 shells and incubated at 70% saturation. Shells

were dried for 4 hours at 45°C and left in a desiccator overnight to cool. The shells were

weighed on an SARTORIUS ultramicro balance with a precision of ±1 µg.

The subsamples were incubated individually in acid leached 1-L PE-bottles. To avoid the

accumulation of shells and the formation of microzones with slightly increased saturations

(“artificial sediment“, Morse et al., 1979), the bottles were fit on a plankton wheel at lowest

spinning velocity (~1-1.5 turns/min). The incubation of all subsamples started at the same time,

while completion of each experiment depended on the different saturation states. Total

incubation time for individual samples thus lasted between 5 and 77 days (Table 5). Upon

termination of the experiment, the foraminifera were rinsed in ultrapure water, dried as described

above, weighed and counted to determine the loss of calcite per shell.

Publication III 46

Sample Preparation and Analytical Techniques

Stable oxygen and carbon isotopes

Stable isotope analyses of the calcite samples at the University of California, Davis,

required ≥15 µg of sample for precise measurements. Single shell analyses on G. sacculifer

resulted in an isotopic range that was much larger than the analytical uncertainty. Due to this

large natural variability we pooled 50 shells, crushed them and analyzed smaller subsamples.

Care was taken to select shells within 50 µm size increments for pooled shell analyses to

minimize ontogenetic factors that affect carbon and oxygen isotopes in foraminifera shells (e.g.

Berger et al., 1978). Duplicate analyses of three such samples of undissolved shells gave

reproducible results within the analytical uncertainty.

Depending on the dissolution state of N. pachyderma (sin.), 2-4 shells were pooled to

yield sufficient carbonate. Between 2 and 6 replicate analyses were conducted for each

dissolution state.

Samples were roasted at 375°C in vacuo for 30 minutes and analyzed with a Fisons

Optima isotope ratio mass spectrometer using an Isocarb common acid bath autocarbonate

system at 90°C. Results are given in delta notation (δ13C = [(13C/12Csample/13C/12Cstd–1)*1000]

and δ18O = [(18O/16Osample/18O/16Ostd–1)*1000]) relative to the Vienna Pee Dee belemnite

standard (V-PDB). The external precision of the measurements is ±0.06‰ and ±0.05‰ for δ18O

and δ13C respectively, based on repeat analyses of NBS 19 standard.

Minor elements

Preparation of foraminifera for trace metal analyses followed the rigorous multistep

cleaning method established by Boyle (1981) and Boyle and Keigwin (1985/1986). As the initial

design of the experiment included Ba/Ca, U/Ca and Cd/Ca, treatment with an alkaline chelating

solution was added according to the procedure by Lea and Boyle (1989) and Lea (1993). Briefly,

50 shells per sample were gently crushed and split in two. Half of each sample was cleaned using

the above named process. In order to test the corrosivity of the cleaning method, the other half

was repeatedly ultrasonicated in methanol and ultrapure water and then transferred to acid

leached microcentrifuge vials. Further cleaning steps were omitted for these splits.

Sample analysis followed the multi-element ICP-MS method previously described by Lea

and Martin (1996). After cleaning, 20-250 µg of purified foraminifera shells were dissolved in

dilute HNO3 solution containing calibrated concentrations of 25Mg, 45Sc, and 89Y. The solutions

Publication III 47

were aspirated into a Finnigan Element 2 high resolution magnetic sector ICP-MS and Mg, Ca

and Sr acquired by analog detector mode. The metal to Ca2+ ratios were estimated by isotope

dilution (25Mg/24Mg) and internal standard calculation (45Sc/48Ca and 89Y/88Sr). Na/Ca was

determined to confirm that the hydrogen peroxide – sodium hydroxide solution used in sample

preparation was completely rinsed out. Several analyses of a consistency standard with trace

metal and Ca2+ concentrations similar to the foraminiferal samples had a standard deviation of

0.83% for Mg/Ca and 0.7% for Sr/Ca (pers. comm. G. Paradis, UC Santa Barbara). Analyses

were performed at the University of California, Santa Barbara.

Additional Mg/Ca analyses were carried out using a wavelength-dispersive, automated

four-spectrometer CAMECA electronmicroprobe. Measurements were done according to the

method described by Nürnberg (1995). Briefly, shells of G. sacculifer (315-450 µm size fraction)

were rinsed in distilled water, embedded in resin, and ground and polished in order to reveal

fresh, relief-free calcite surfaces of chamber walls. The electron beam was focussed on spots

approximately 2-4 µm in diameter. Element concentrations in the ppm-range were detected

quantitatively, calculated stoichiometrically in oxide form, and reported as element/Ca ratios.

Mg-Periclase (synthetic) and Ca-Wollastonite (natural) served as standards. Between 26 and 54

spot analyses have been carried out on each of 10 individual shells. Measurements were done at

GEOMAR in Kiel, Germany.

Stable boron isotopes

Shells were crushed and bleached with 4-6% sodium hypochlorite to remove organic

matter. After at least 24 hours in bleach, the samples were rinsed 10 times in ultrapure water,

ultrasonicated and centrifuged to rinse out the bleach and remove soluble salts and adsorbed

boron. The cleaned carbonate was dissolved in 2N quartz distilled HCl, loaded on a rhenium

zone refined filament, and 1 µl of boron-free seawater was added to enhance ionization and

suppress fractionation (Hemming and Hanson, 1994). Isotope data were collected on a Finnigan

MAT 262 RPQ+ Thermal Ionization Mass Spectrometer (TIMS) at GEOMAR in Kiel, Germany.

The BO2- ion method followed the procedure outlined in Sanyal et al. (1996) and Sanyal et al.

(1997). Each sample was run at least 4 times. Analysis temperature was constant at 940 ± 10°C.

Publication III 48

Table 5. Average weights of undissolved and dissolved foraminifera shells, dissolution estimates and calcite saturation of experimental seawater.

target saturation and incubation time

index

initial saturation

state incubation time initial

weight (µg/shell)

final weight

(µg/shell)

% weight

loss

number of shells

recovered (d)

(Ω) (t1-t5)

G. sacculifer, 450-500 µm, 35 shells per sample 120% t1 1.18 14 36.4 33.1 9.1 32 t2 29 37.7 33.9 9.9 32 t3 48 39.2 38.7 1.2 34 t4 63 40.2 39.1 2.8 34

t5 77 40.3 37.2 7.7 32 90% t1 0.88 9 40.9 40.2 1.5 35 t2 22 40.9 40.5 1.0 33 t3 34 41.2 36.1 12.5 31 t4 55 41.9 41.7 0.4 34 t5 63 42.7 36.7 14.1 31 70% t1 0.69 7 42.9 39.8 7.4 34 t2 14 43.7 39.8 8.8 34 t3 22 43.9 39.7 9.6 34 t4 34 43.9 32.6 25.8 33 t5 55 44.9 31.1 30.7 32 50% t1 0.49 5 45.4 34.3 24.5 30 t2 9 45.9 35.2 23.3 32 t3 14 46.1 32.7 29.0 35 t4 19 47.5 26.2 44.9 32 t5 22 48.3 27.5 43.0 35 G. sacculifer, 315-450 µm, 100 shells (* = 500 shells) per sample 120% t1 1.24 14 25.4 24.1 4.9 96 t2 28 26.3 25.6 2.6 98 t3 43 26.6 25.4 4.6 98 t4 56 27.5 26.7 2.7 95 t5 70 27.3 22.9 16.1 94 90% t1 0.84 14 27.9 26.4 5.4 99 t2 28 28.3 22.8 19.6 85 t3 43 28.9 17.2 40.5 89 t4 56 29.2 19.7 32.6 86 t5 70 29.4 19.4 34.0 93 70% t1 0.69 8 29.4 26.1 11.3 96 t2 13 29.6 23.5 20.9 97 t3 17 30.2 23.1 23.5 96 t4 20 30.2 20.7 26.9 91 t5* 24 28.3 23.0 18.5 484 50% t1 0.56 8 31.1 25.6 17.8 99 t2 13 31.9 20.1 37.0 91 t3 17 32.5 15.0 53.8 86 t4 20 33.1 14.7 55.7 89 t5* 24 28.5 14.5 49.0 404 N. pachyderma (sin.), 279-309 µm, 100 shells per sample 70% t1 0.75 6 10.9 10.8 1.3 94 t2 14 11.2 10.6 5.8 93 t3 28 11.2 9.3 17.4 96 t4 40 11.3 9.0 20.6 98 t5 49 11.3 5.8 48.3 90 t6 63 11.4 3.4 69.9 39

Publication III 49

To rule out isobaric interferences on mass 42 due to residual organic matter

(12C14N16O-ions), we checked for mass 26 (12C14N-ions) during each measurement. No

interferences were detected. The 11B/10B ratio was corrected for isotopic interferences on mass

43 (10B16O17O-ions) by subtraction of 0.00078 from the 43/42 ratio (Spivack and Edmond,

1986). Results are reported in delta notation (δ11B=[(11B/10Bsample/11B/10Bstd–1) * 1000]

relative to the NBS 951 boric acid standard. Repeated analyses of natural seawater used as a

laboratory standard resulted in an average value of 39.5 ± 0.3‰.

Determination of δ44Ca

For N. pachyderma (sin.), 2-4 shells per dissolution state were selected and pooled.

Due to the strong seasonal signal in δ18O, the number of individuals of G. sacculifer (450-500

µm size fraction) was raised to 15. Three samples of each, undissolved and dissolved G.

sacculifer shells, were prepared for isotope analysis. Sample analysis followed the method

established by Nägler et al. (2000). Briefly, samples were dissolved in 2N ultrapure HCl and

spiked with a 43Ca-48Ca double spike. Samples were loaded on rhenium filaments and

measured on the Finnigan MAT 262 RPQ+ TIMS at GEOMAR. Each sample was run 1-3

times. Data are presented in delta notation relative to a natural CaF2 standard:

(δ44Ca=[(44Ca/40Casample/44Ca/40Ca std–1)*1000]. External reproducibility is 0.1‰.

Results Dissolution state and solubility

The respective dissolution state of each sample is reported in terms of weight loss [%]

per shell. Weight loss data and saturation states of the experimental seawater are shown in

Table 5. The calcite dissolution rates obtained in this study are in good agreement with

previous research and show both, intra- and interspecific variability in solution susceptibility.

The larger size fraction of G. sacculifer proved more solution resistant than the smaller size

fraction and G. sacculifer was generally more prone to dissolution than N. pachyderma (sin.).

Similar results have been obtained by Berger (1967; 1970) in studies of differential

dissolution on foraminiferal assemblages within the natural environment.

Visual inspection of the shells under the light microscope at 66-fold magnification

does not allow the detection and quantification of partial dissolution. Already at 5-6% weight

loss shells loose their shiny appearance and get opaque. Until disintegration starts and the

texture of severely dissolved shells gets almost sponge-like, wall microstructural changes

Publication III 50

a b

c d Figure 12. Microstructural breakdown of G. sacculifer shell surfaces monitored by scanning

electron microscopy (magnification: 800x). (a) Undissolved shell with broad interpore ridges and filled pore openings as typical for gametogenic specimens. (b) Average shell weight loss ~22%, first fissures appear inside of the pore funnels and spine holes deepen. (c) ~ 26% dissolved, fissures also cut through the interpore ridges and spine bases. (d) ~44% dissolved, deep cracks cover the entire shell surface and interpore ridges are remarkably rounded.

cannot be detected under the light microscope. Scanning electron microscopy allows a much

more detailed examination. Since Helmke and Bauch (2002) have presented similar

investigations for N. pachyderma (sin.), we focus here on G. sacculifer. Figure 12 shows

different preservation states of the outer shell structure. Figure 12a is a typical picture of a

gametogenic shell surface as described by Bé (1980) with broad interpore ridges, covered

spine stubs, and filled pore openings. The shell was incubated for 77 days at 120% calcite

saturation and no indication of dissolution is visible. Figure 12b shows the onset of

dissolution with first fissures inside of the pore funnels and deepening of the so called “spine

holes“, small craterlike depressions at the former position of the spines during ontogenetic

Publication III 51

shell growth (Bé, 1980). Upon further dissolution the fissures also cut through the interpore

ridges and spine bases (Fig. 12c) until at 44% weight loss deep cracks cover the entire shell

surface (Fig. 12d). This observation demonstrates that shell corrosion leaves a systematic

fingerprint on partially dissolved shells which can be easily detected by SEM examination.

Minor elements: Mg/Ca and Sr/Ca

Bulk shell chemistry data as determined by ICP-MS are presented in Figures 13 and

15. Analytical precision of consistency standards is much better than the reproducibility of

biogenic samples. Lea et al. (1999b) stated that high Na/Ca indicates inadequate rinse-out of

leach and that samples with Na/Ca >8 µmol/mol may be contaminated. In our study, rejection

of data with Na/Ca >8 µmol/mol removed outliers. However, Na/Ca cutoffs also result in a

high rejection rate and the likely removal of uncontaminated data. Comparison of the two

cleaning procedures demonstrates that the 8 µmol/mol level is sometimes exceeded in

samples that were merely rinsed and did not see any leach during cleaning. Therefore, only

the most obvious outliers were rejected from the thoroughly cleaned samples (Table 6).

2.5

3.0

3.5

4.0

4.5

5.0

5.5

-10 0 10 20 30 40 5

G. sacculifer

Mg/

Ca

(mm

ol/m

ol)

dissolution (% weight loss)

a

00.4

0.5

0.6

0.7

0.8

-10 0 10 20 30 40 50

N. pachyderma (sin.)

Mg/

Ca

(mm

ol/m

ol)


b

Figure 13. Mg/Ca in G. sacculifer (a) and N. pachyderma (sin.) (b). Symbols in (b) indicate

different analytical pretreatments: Samples were either rinsed in methanol and ultrapure water (open squares) or thoroughly cleaned according to methods established by Boyle (1981), Boyle and Keigwin (1985/1986) and Lea (1993) (closed triangles). Merely rinsed samples confirm the trend observed in thoroughly cleaned shells. Open circles in (a) indicate samples that were actually too large for the calibrated analytical size range and element ratios thus may be underestimated. Note that G. sacculifer Mg/Ca remains constant despite a more than 5 times higher Mg concentration than in N. pachyderma (sin.).

Publication III 52

Table 6: Experimental data: Minor and trace elements.

sample calcite µg

Mg/Ca mmol/mol

Sr/Ca mmol/mol

Na/Ca mmol/mol

G. sacculifer, 315-450 µm rinse only

undissolved 485 5.29 1.37 7.08undissolved 389 4.97 1.36 7.06undissolved 445 4.75 1.36 6.94

120% t3 490 4.14 1.36 7.13120% t4 576 3.90 1.34 6.94120% t5 234 4.47 1.35 7.19

90% t1 462 4.47 1.35 6.9790% t2 258 4.30 1.36 7.0690% t3 31 4.14 1.39 9.4490% t4 193 3.99 1.38 7.3670% t1 372 4.41 1.37 7.0470% t2 204 3.88 1.36 7.4870% t3 169 3.82 1.36 6.9170% t4 135 3.74 1.37 7.1850% t1 296 4.01 1.38 7.2150% t2 151 4.12 1.36 7.6450% t3 113 10.69* 1.37 7.5350% t4 56 4.74 1.37 8.43

rinse, reduce, leach undissolved 216 4.12 1.35 7.13undissolved 132 4.21 1.37 6.78undissolved 223 4.14 1.38 6.77

120% t3 454 3.79 1.37 6.53120% t4 301 3.90 1.36 6.76120% t5 434 3.72 1.36 6.64

90% t1 371 3.87 1.37 6.6390% t2 334 3.90 1.37 6.5090% t3 28 4.05 1.39 8.6290% t4 89 3.84 1.38 6.6670% t1 418 4.18 1.37 6.5070% t2 257 3.74 1.37 6.8770% t3 240 3.61 1.38 6.7870% t4 192 4.06 1.38 7.2050% t1 323 3.87 1.39 7.4950% t2 103 3.88 1.39 6.94

N. pachyderma (sin.) rinse only

70% t1 122 0.70 1.54 7.3670% t2 143 0.62 1.55 7.3270% t3 98 0.56 1.57 7.7270% t4 21 0.48 1.59 10.1370% t5 16 0.83* 1.60 10.91

rinse, reduce, oxidize, chelate, leach 70% t1 171 0.74 1.57 7.3170% t2 89 0.69 1.58 7.9870% t3 38 0.60 1.62 10.1070% t4 63 1.21* 1.64 9.03

Samples are splits of 50 crushed shells. Splits were either merely rinsed in methanol and ultrapure water or chemically pretreated according to the method established by Boyle (1981), Boyle and Keigwin (1985/1986) and Lea (1993). See text for details. Each value represents a single analysis. Sample identification according to Table 1. For G. sacculifer, samples with >250 µg calcite exceed the analytically calibrated sample size. Element ratios of these samples may be underestimated. Na/Ca >8mmol/mol may indicate inadequate rinse-out of bleach. * = samples omitted from figures due to likely contamination (see text).

Publication III 53

The Mg/Ca ratio in shells of N. pachyderma (sin.) is clearly reduced upon dissolution

(Fig. 13b). Applying the sediment core-top calibration of Nürnberg (1995) for left coiling N.

pachyderma (T°C = ln (Mg/Ca / 0.55) / 0.099), a temperature estimate of 3°C for least

dissolved shells and 0.9°C for shells that lost 17% of their initial weight can be calculated.

The temperature estimate is thus lowered by ~2 K.

Some samples were accidentally heavier than the weight range for which the method

was calibrated. The element ratios of these samples may be underestimated and we highlight

them in all plots as open symbols. However, a trend towards underestimated values was not

observed for any of the elements investigated. Data on N. pachyderma (sin.) are not affected

by this problem since all samples were within the calibrated weight range.

For G. sacculifer, no significant dissolution effect on Mg/Ca can be observed (Fig.

13a). The average Mg/Ca is 3.93 ± 0.18 mmol/mol. For comparison, microprobe analyses on

10 G. sacculifer shells yield clearly heterogeneous Mg/Ca distributions in individual shells.

Ratios for outer shell calcite average 4.61 ± 0.47 mmol/mol and 3.18 ± 0.54 mmol/mol for

inner calcite (Fig. 14). Since microprobe analyses are carried out without elaborate

preparation and cleaning procedures, it could be argued that the enrichment is due to

contaminating phases. A simple mass balance calculation based on the assumption that G.

sacculifer adds ~28% (Bé, 1980) of its shell weight during gametogenesis, yields a bulk shell

Mg/Ca ratio of 3.58 ± 0.52 mmol/mol. Given the large seasonal temperature variability at the

core site and the low number of shells investigated using microprobe compared to ICP-MS

measurements, the similarity between the two approaches is promising.

0 1 2 3 4 5 6 7

123456789

10G. sacculifer

Mg/Ca (mmol/mol)

shel

l No.

Figure 14. Averages of Mg/Ca in inner (grey bars) and outer calcite (black bars) of 10 G. sacculifer shells as determined by microprobe analysis of wall profiles. Each value is the average of 15-28 spot analyses. Approximately 20% higher Mg concentrations are encountered in outer compared to inner calcite.

Publication III 54

1.52

1.56

1.60

1.64

-10 0 10 20 30 40 50


Sr/C

a (m

mol

/mol

)


b

1.34

1.36

1.38

1.40

-10 0 10 20 30 40 50

G. sacculiferSr

/Ca

(mm

ol/m

ol)


a

Figure 15. Sr/Ca in G. sacculifer (a) and N. pachyderma (sin.) (b). Both foraminifera species

show a slight increase upon dissolution. Symbols as in Figure 13.

Sr/Ca ratios increase for both species upon dissolution. The observed effect is smaller

in G. sacculifer where the increase from 1.36 to 1.39 mmol/mol (at ~40% weight loss) is

equal to 2.2% (Fig. 15a). Values in shells of N. pachyderma (sin.) increase by 4.5% from 1.57

to 1.64 mmol/mol at merely 20% weight loss (Fig. 15b). Complete cleaning procedures seem

to further increase the Sr/Ca ratio.

Stable oxygen and carbon isotopes

Data are presented in Figures 16 and 17. Although repeated measurements of 50

pooled shells of G. sacculifer gave results within the uncertainty of the analysis, scatter of the

final dissolution data is still large. Results average at δ18O = –0.22 ± 0.31 ‰ (V-PDB) and

δ13C = 2.18 ± 0.38 ‰ (V-PDB). For none of the two isotopic compositions significant

changes occur upon dissolution.

Although data are overall very variable, carbon and oxygen isotopic composition tend

to increase in shells of N. pachyderma (sin.) upon dissolution (Fig. 16b, 17b). We do not have

a δ18O-temperature relationship for left coiling N. pachyderma to determine the temperature

underestimate of dissolved shells. Applying the slope of von Langen (2001) for the right

coiling type (∆T (K) = – 6.16 * δ18Ocalcite (V-PDB)) yields a maximum difference of –1.6 ± 1.2 K

between the least and the most dissolved shells.

Publication III 55

-0.5

-0.4

-0.3

-0.2

-0.1

0.0

0.1

-10 0 10 20 30 40 50 60

G. sacculiferδ18

O (‰

PD

B)


a

3.6

3.8

4.0

4.2

4.4

4.6

0 20 40 60 80


δ18O

(‰ P

DB

)


b

Figure 16. δ18O values in G. sacculifer (a) are averages of duplicate analyses on samples of 50 pooled shells. In agreement with the lack of thermal and salinity gradients in the water column, G. sacculifer δ18O is suggested to be distributed homogeneously throughout the shell and does not change with dissolution. (b) N. pachyderma (sin.). The increase in δ18O converts to a temperature underestimate of -1.6 ± 1.2 K for the most dissolved shells. Data are averages of 2-6 replicate samples with 2-4 shells each. Error bars are standard errors of multiple samples.

1.7

1.8

1.9

2.0

2.1

2.2

2.3

-10 0 10 20 30 40 50 60

G. sacculifer

δ13C

(‰ P

DB

)


a

-0.2

0.0

0.2

0.4

0 20 40 60 80


δ13C

(‰ P

DB

)


b

Figure 17. δ13C in G. sacculifer remains constant (a) and increases slightly in N. pachyderma (sin.) (b). Symbols and error bars as in Figure 16.

Stable calcium and boron isotopes

Within three samples of 15 undissolved G. sacculifer shells variability in δ44Ca is as

large as the variability between shells that have been dissolved to different degrees (Fig. 18a).

With regard to N. pachyderma (sin.), a slight increase in the isotope value occurs upon

continuing shell weight loss (Fig. 18b). The observed change barely exceeds the external

reproducibility of the analysis.

Publication III 56

-1

-1

-0.8

-0.6

-0.4

-0.2

0 10 20 30

G. sacculiferδ44

Ca

(‰)


a

-1.0

-0.9

-0.8

-0.7

-0.6

0 20 40 60 80



δ44C

a (‰

)

b

Figure 18. δ44Ca is based on replicate analyses of samples with 15 large G. sacculifer shells

((a) size fraction 450-500 µm) and 3 N. pachyderma (sin.) shells (b). Error bars are standard deviations of multiple samples runs. External reproducibility of the analysis is 0.1‰.

The δ11B in shells of G. sacculifer increases by approximately 1.2‰ between undissolved

shells and shells that have lost 50% of their initial weight (Fig. 19). In addition to the isotopic

increase, signal intensity during analysis was observed to be reduced in the dissolved samples.

Since the volume of acid used to dissolve a sample was adjusted to the amount of calcite,

reduced signal intensity most likely indicates an overall lowered boron concentration in the

dissolved samples.

20

21

22

23

-10 0 10 20 30 40 50 60

G. sacculifer

δ11B

(‰)


Figure 19. δ11B in G. sacculifer is shifted towards heavier values with increasing dissolution.

Closed circles are averages of 4 replicate analyses on 400-500 pooled shells. Errors are expressed as 2σmean for multiple sample analyses.

Publication III 57

Discussion A conceptual model

One of the hypotheses that have been proposed to explain changes in foraminiferal

shell chemistry is the preferential dissolution of thinshelled individuals which are supposed to

have a different shell chemistry than thicker shelled individuals (Boyle and Rosenthal, 1996;

Erez, 1979a). Although we cannot rule out that non-encrusted individuals may be present in

sediments at other locations, all 25 shells examined by SEM for this study clearly showed

indices of encrustation (see also Figure 12). Furthermore, shell recovery in our experiments

was high (Table 5) and the few lost shells are due to random loss during sample handling and

fragmentation due to mechanical damage rather than selective dissolution. The chemical

changes observed in this study therefore have to be attributed to chemical changes within

individual shells.

Previous studies have already pointed out that foraminifera secrete different structural

types of calcite during their lifecycle and that chemical and isotopic differences exist between

these calcites. In particular, some species are known to deposit a thick crust during late

ontogeny which is usually associated with gametogenesis and often occurs in deeper waters of

different physico-chemical conditions (e.g Bé, 1980; Caron et al., 1990; Duplessy et al.,

1981). In G. sacculifer, gametogenic calcite is deposited over a period of only 16 hours (Bé,

1980). Considering that ontogenetic chamber formation in G. sacculifer takes ~4 weeks

(Bijma et al., 1994) and contributes ~2/3 to the final shell weight, the precipitation rate for

gametogenic calcification is very high and leads to structural differences between inner and

outer calcite (Bé, 1980).

In the literature, naming of the two calcite types is very confusing. The proximal type

may be described as inner, primary or ontogenetic calcite and the distal as outer, secondary,

gametogenic or crust calcite. With regard to partial shell dissolution in the sediment it was

suggested that ontogenetic calcite dissolves faster than the gametogenic crust which would

lead to a systematic shift of the initial bulk shell chemistry to the chemistry of the outer

calcite (Brown and Elderfield, 1996; Dekens et al., 2002; Lohmann, 1995; Rosenthal et al.,

2000; Russell et al., 1994). We use this concept to examine the dissolution behavior of the

shells studied herein.

Figure 20 illustrates the simplified two component system of a G. sacculifer that

secretes ~2/3 of its shell close to the sea surface and adds ~28% gametogenic calcite (Bé,

1980) at depth. The chemical and isotope values assigned to inner and outer calcite are based

on the bulk chemistry measured on our G. sacculifer shells (Figure 20). Mg/Ca for inner and

Publication III 58

outer calcite is known from microprobe analyses but the other values need to be assumed

based on theoretical considerations. Hydrographic data collected at the core site report that

vertical temperature and salinity differences are smaller than 0.2 K and 0.1‰ over the entire

depth range (Pätzold and participants, 2000). Upon migration to deeper waters in late

ontogeny of G. sacculifer, individuals are therefore supposed to have encountered

homogeneous temperatures and salinities but lower pH (derived from the nearby GEOSECS

station 405). While Erez et al. (1991) argue that gamete release of G. sacculifer in the Gulf of

Elat may occur at a depth of 200 m or even below, Bijma and Hemleben (1994) suggest that

G. sacculifer reproduces at ~80 m water depth in the central Red Sea. Because we do not want

to underestimate possible chemical differences between inner and outer calcite, we assume the

physico-chemical conditions at 250 m water depth for estimating the chemistry of the

gametogenic calcite. Based on these theoretical reflections we would expect lower δ11B in the

gametogenic calcite but similar δ18O, δ44Ca and Sr/Ca in inner and outer calcite (Sr

distribution in the water column is conservative). Mg/Ca is the only parameter which relies

directly on measured data (cf. Figure 14). δ13C is excluded from this model because in

addition to the "carbonate ion (or pH) effect" (Bijma et al., 1999; Spero et al., 1997) the signal

recorded in foraminiferal shells also depends on vertical changes in primary productivity

(changes in the seawater isotopic composition). Due to the lack of seawater isotope values we

cannot quantify these effects for this locality.

The higher Mg/Ca measured in outer compared to inner calcite is counterintuitive as

temperature is known to be the dominant parameter affecting Mg incorporation in

foraminiferal shells (Lea et al., 1999b). Hence, under constant water temperature conditions

we would expect similar Mg/Ca in inner and outer calcite. However, Nürnberg et al. (1996)

observed a similar pattern in laboratory grown G. sacculifer. They speculated that some sort

of a vital effect was responsible for the observed Mg variation and that this vital effect would

be usually hidden by the dominant temperature influence experienced by naturally migrating

foraminifera. Recently, Elderfield et al. (2002) observed that Mg/Ca increases with increasing

shell size. Because δ18O in the same shells does not follow this pattern Elderfield et al. (2002)

argued that temperature could not explain the increase but that Mg uptake must be a

nonequilibrium process which is partly decoupled from the dominant temperature effect. Our

shells are similar to Nürnberg et al.'s laboratory grown individuals in that they did not

experience temperature changes.

Publication III 59

ontogenetic (inner) calcite Mg/Ca = 3.18 mmol/mol Sr/Ca - 1.36 mmol/mol δ18 O - -0.28 ‰ (V-PDB) δ11 B - 20.9 ‰ at pH=8.25 δ44 Ca - -0.62 ‰

gametogenic (outer) calcite Mg/Ca = 4.61 mmol/mol Sr/Ca - 1.36 mmol/mol δ18 O - -0.28 ‰ (V-PDB) δ11 B - 19.9 ‰ at pH=8.12 δ44 Ca - -0.62 ‰

measured bulk shell chemistry of undissolved shells Mg/Ca = 4.16 ± 0.05 mmol/mol Sr/Ca - 1.36 ± 0.02 mmol/mol δ18 O - -0.28 ± 0.1 ‰ (V-PDB) δ13 C - 1.99 ± 0.06 ‰ (V-PDB) δ11 B - 20.77 ± 0.28 ‰ δ44 Ca - -0.62 ± 0.12 ‰

sedimentation

ontogenetic migration

gametogenic calcification at ~250m

ontogenetic calcification in surface water

Figure 20. Schematic presentation of vertical migration and varying calcification depths. The model G. sacculifer is assumed to secrete ~2/3 of the bulk calcite at the surface and ~28% as a gametogenic crust (Bé, 1980) at ~250 m depth (Erez et al., 1991). The model assumptions do not account for the juveniles' ascend from and subadults' descend to the reproduction depth. Chemical composition of ontogenetic and gametogenic calcite were either directly measured (Mg/Ca, cf. Figure 14) or estimated according to theoretical reflections on the physico-chemical gradients encountered at the core site (δ18O, δ11B, δ44Ca, Sr/Ca). A mass balance calculation was applied to bring theoretical reflections and measured bulk shell chemistry data into line.

Based on previous suggestions that shell chemistry is determined by a faster

dissolution rate of ontogenetic relative to gametogenic calcite, the following dissolution

pattern would be expected: no change in δ18O, δ44Ca and Sr/Ca, an increase in Mg/Ca and a

decrease in δ11B. With regard to Mg/Ca the story is a little bit more complicated: The Mg

content of biogenic calcite has been proposed to promote dissolution (Brown and Elderfield,

1996). Because higher Mg-incorporation usually predominates during early ontogeny in

warmer surface waters, the theory of higher solution susceptibility due to higher Mg content

Publication III 60

usually coincides with dissolution of early ontogenetic calcite. In this special case (higher

Mg/Ca in outer calcite) Mg/Ca should increase if the inner calcite dissolved faster than the

outer calcite or decrease if the higher Mg content of the outer calcite was sufficient to increase

the dissolution susceptibility of the outer calcite. On this background we will now discuss the

results obtained in this study.

Minor elements

The observed Mg/Ca decrease in shells of N. pachyderma (sin.) on the order of 20-

30% at only 17% weight loss is sizable and translates to a temperature underestimate of at

least 2 K in these shells. Neogloboquadrina pachyderma (sin.), migrates from subsurface

waters to 50-200 m water depth and this change of habitat is reflected in increasing δ18O of

this species (Kohfeld et al., 1996). Due to the encountered change in temperature it is likely

that Mg/Ca in N. pachyderma (sin.) displays a similar pattern to δ18O so that the observed

decrease in Mg/Ca is in line with the suggestion that ontogenetic calcite is preferentially

dissolved due to its higher magnesium content (Brown and Elderfield, 1996; Lorens et al.,

1977; Savin and Douglas, 1973).

For G. sacculifer a dissolution effect similar to the one in N. pachyderma (sin.) was

not observed despite the heterogeneous Mg distribution. This result is neither consistent with

a higher dissolution susceptibility of ontogenetic calcite (which would predict increasing

Mg/Ca for our shells) nor with Brown and Elderfields model that calcite with a higher Mg

content is less stable (which would predict decreasing Mg/Ca). As the difference in Mg/Ca

between outer (4.61 mmol/mol) and inner calcite (3.18 mmol/mol) is not as large as the one

determined for Globorotalia tumida (2.1 and 5.1 mmol/mol, respectively, Brown and

Elderfield, 1996), it may well be that it is too small to result in a measurable dissolution

effect. To test this hypothesis, we included the ion activity product (IAP) equation established

by Brown and Elderfield (1996) into a mass balance calculation and estimated the Mg/Ca of

shells that have lost 50% of their initial weight:

IAP = 4.17 * 10-9 + 1.97 * 10-11 * Mg/Ca (10) dissolution rate = κ * (1-Ω)η (11)

Ω = a

Ca 2+ * aCO32−

IAP (12)

where κ (d-1) denotes the dissolution rate constant, η (dimensionless) the dissolution reaction

order and Ω the saturation state with respect to Mg-calcite. As the rate constant κ is different

Publication III 61

for different biogenic carbonates and the reaction order η is not very well constrained, we

used several values cited in the literature (Hales and Emerson, 1997b; Keir, 1980; Morse,

1978; Walter and Morse, 1985). aCa2+ and aCO32- are the activities of Ca2+ and CO32- in

seawater which can be calculated using the seawater concentrations of [Ca2+] = 10.3 mmol kg-

1 and [CO32-] = 35 µmol kg-1 (value chosen in accordance to our experimental seawater

carbonate chemistry at 90% saturation, 15°C and atmospheric pressure) and the total activity

coefficients of these ions in seawater as provided by Millero and Pierrot (1998) (γtCa2+ = 0.203

and γtCO32- = 0.039). The amount of crust calcite secreted by G. sacculifer and G. tumida is

assumed to account for ~28% (Bé, 1980), respectively ~70% (Brown and Elderfield, 1996) of

the undissolved shell weight. Due to the heterogeneous Mg distribution, the relative

proportions of inner and outer calcite are supposed to change systematically with increasing

dissolution and so should the bulk Mg/Ca.

The results of our calculations are shown in Table 7. For any κ and η chosen, no

significant change in Mg/Ca of G. sacculifer can be determined. Similarly, 50% weight loss

in G. tumida does not reproduce the Mg/Ca variation (i.e. from 2.65 to 1.25 mmol/mol)

observed by Brown and Elderfield (1996). We also tested a much larger heterogeneity

between inner and outer calcite (10 and 2 mmol/mol, respectively). Only such a large contrast

between inner and outer calcite gives rise to a significant Mg/Ca reduction from 7.76 to 7.27

mmol/mol. However, using calibration equations for the temperature dependency of Mg/Ca in

G. sacculifer (Dekens et al., 2002; Nürnberg et al., 1996), such a heterogeneity would require

temperature differences of >17 K between surface habitat and encrustation depth.

We thus have to conclude that the difference in the Mg content between inner and

outer calcite is too small to cause a detectable dissolution effect. In addition, the strong Mg

dissolution effect observed in N. pachyderma (sin.) is striking with regard to the combination

of low Mg-content (~0.7 mmol/mol) and lower dissolution susceptibility of these shells. All in

all, these calculated dissolution rates argue against the suggestion that higher Mg

concentrations significantly increase the dissolution susceptibility of particular shell parts. We

conclude that Mg is either not the only dissolution-driving factor or that the Mg content is

coupled to some other parameter which is responsible for the overall dissolution behavior. An

alternative (or additional) explanation for different dissolution susceptibilities of inner and

outer calcite could be differences in surface area and/or crystal structure. A recent review

article by Morse and Arvidson (2002) summarizes that complex microstructure and reactvity

of the surface area of biogenic magnesian calcites strongly control the solubility. With regard

to G. sacculifer, the slowly secreted ontogenetic calcite shows well-developed polygonal

Publication III 62

patterns, whereas gametogenic calcite is deposited within a few hours and the crystal

structure is rather irregular (Bé, 1980). On this background, one could imagine that the

structural differences affect the dissolution kinetics of the two calcites so that the use of the

same dissolution rate constant κ for both shell components as done for the dissolution rate

calculations (Table 7) may not be appropriate.

Table 7: Numerical experiment on the dissolution susceptibility of foraminiferal Mg-calcites

Mg/Ca

mmol/mol

bulk

Mg/Ca mmol/mol

IAP

Ω

dissolution

rate4

bulk Mg/Ca (mmol/mol)

50% dissolved

undissolved Keir (1980)a

Walter & Morse (1985)b

Morse (1978)c

Hales & Emerson (1997)d

G. sacculifer (this study)1

outer 4.61 3.58 4.26*10-9 0.670 0.0341 3.56 3.57 3.57 3.58 inner 3.18 4.23*10-9 0.674 0.0321 theoretically2

outer 2.0 7.76 4.21*10-9 0.678 0.0305 7.27 7.43 7.42 7.64 inner 10.0 4.37*10-9 0.654 0.0424 G. tumida (Brown & Elderfield, 1996)3

outer 2.1 2.85 4.21*10-9 0.678 0.0306 2.77 2.80 2.80 2.83 inner 5.1 4.27*10-9 0.668 0.0348

Mg/Ca ratios for inner and outer shell calcite were determined by electron microprobe analysis (1cf. Figure 14; 3Brown and Elderfield, 1996) and are averages over a large range of measured values. Bulk Mg/Ca of the undissolved shells was calculated assuming that 1,2 28%, respectively 3 70% of the undissolved shell calcite are crust calcite (Bé, 1980; Brown and Elderfield, 1996). Due to the heterogeneous Mg distribution, dissolution systematically changes the relative proportions of inner and outer calcite. The corresponding changes in bulk Mg/Ca were calculated for shells that have lost 50% of their initial weight due to dissolution. Parameterization of dissolution kinetics according to text equations (1) - (3). Parameter values are a κ=5, η=4.5; b κ=1, η=2.9; c κ=5, η=3; d κ=0.38, η=1. Ω was calculated for seawater-[CO3

2-] = 35 µmol kg-1 and atmospheric pressure. 4 Presented dissolution rates were calculated according to parameter values of Keir (1980) and are meant to give an impression of the small rate differences between foraminiferal shells with different Mg contents. Dissolution rates vary with the values for κ and η chosen.

In order to determine how much faster one shell component has to dissolve relative to

the other to result in a measurable change of a heterogeneously distributed element, we

calculated the change in Mg/Ca in response to varying dissolution rates of gametogenic and

ontogenetic calcite (Figure 21). For these calculations we used the measured Mg/Ca values

for our G. sacculifer shells and assumed that gametogenic calcite amounts to ~28% of the

Publication III 63

undissolved shell (Bé, 1980). Although it was said before that the higher Mg content of the

outer shell calcite is unusual for foraminifera that experience vertical temperature gradients,

the plot would only be mirrored at the 1:1-dissolution ratio axis (Fig. 21) for foraminifera

containing more Mg in the inner shell calcite. Hence, irrespective of which component is

enriched in Mg/Ca, dissolution rates have to be several times higher in one component

relative to the other to result in a change of Mg/Ca on the order of 0.2-0.3 mmol/mol (at 50%

weight loss) which has been observed for Globigerinidae in various sediment studies (e.g.

Dekens et al., 2002; Rosenthal et al., 2000). Having shown that the higher dissolution

susceptibility of Mg-calcite according to the Brown and Elderfield model at small differences

between inner and outer calcite is relatively insignificant for bulk shell Mg/Ca changes, Fig.

21 demonstrates that differences in dissolution rate have to be quite large to result in

measurable effects.

3.2

3.4

3.6

3.8

4.0

4.2

4.4

4.6

0 20 40 60 80 10

ontogenetic vs. gametogenic dissolution

bulk

Mg/

Ca

(mm

ol/m

ol)


1:20

1:4

1:1

4:1

20:1

0

Figure 21. Comparison of different dissolution rates of ontogenetic and gametogenic calcite and the

respective effect on changes of a heterogeneously distributed element in foraminiferal calcite. As a case study, measured Mg/Ca values of G. sacculifer shells were chosen, i.e. 3.18 mmol/mol for ontogenetic and 4.61 mmol/mol for gametogenic calcite. Dissolution rates are relative and should be read as follows: 1:20 means that per dissolution step 1% of gametogenic calcite and 20% of ontogenetic calcite are removed. Equal dissolution rates are assumed at 1:1, at 4:1 4% of gametogenic calcite are dissolved relative to 1% of ontogenetic calcite etc. The relative rate of change in the bulk shell chemistry is determined by the initial amount of gametogenic calcite and therefore may differ between species. We chose 28% according to Bé (1980) for G. sacculifer. At higher gametogenic dissolution rates bulk shell chemistry changes much faster (lower part of the panel) than at higher ontogenetic dissolution rates (upper part of the panel). Finite limits of chemical change are given by the chemistry of outer and inner calcite, respectively.

Publication III 64

Although a few studies have shown that G. sacculifer Mg/Ca may be somewhat less

sensitive to dissolution than other species from the same location (Brown and Elderfield,

1996; Dekens et al., 2002; Hastings et al., 1998), varying degrees of Mg dissolution in G.

sacculifer have been observed in other studies (Dekens et al., 2002; Rosenthal and Lohmann,

2002, in press; Rosenthal et al., 2000). Hence, one might speculate whether the extent of G.

sacculifer's ontogenetic vertical migration changes from site to site so that this species

encounters different physico-chemical gradients at different locations. The G. sacculifer shells

investigated in this study are rather exceptional with their higher Mg/Ca in outer calcite.

While Mg is usually more concentrated in the inner calcite, combination of this pattern with a

lower structural stability of ontogenetic calcite may explain why dissolution drives Mg/Ca to

lower values in shells of foraminifera that encounter vertical temperature gradients during

their life cycle. This could explain why we observe lower Mg/Ca in partially dissolved shells

of N. pachyderma (sin.) whereas in G. sacculifer the structural and Mg-effects may be

balanced and Mg/Ca therefore remains constant.

In situ sediment studies showed that the Sr/Ca ratio of the planktonic G. tumida and

the benthic Cibicidoides wuellerstorffi is lowered by partial dissolution (Brown and

Elderfield, 1996; McCorkle et al., 1995). Brown and Elderfield (1996) concluded that

dissolution initially removes the most impure calcite before dissolving calcite with lower Mg

and Sr concentrations. Similarly, Haley and Klinkhammer (2002) studied the effect of

cleaning and dissolution on Orbulina universa in a flow-through system and observed that Sr

and Mg were removed preferentially. In contrast, shells of G. sacculifer either revealed a

slight increase in Sr/Ca upon dissolution (Bender et al., 1975; Haley and Klinkhammer, 2002)

or no systematic effect at all (Brown and Elderfield, 1996; Lorens et al., 1977). Our

increasing Sr/Ca data on G. sacculifer and N. pachyderma (sin.) line up with the results of

Bender et al. (1975) and Haley and Klinkhammer (2002). The magnitude of change found

herein for G. sacculifer is comparable with changes reported in the literature. For N.

pachyderma (sin.) no dissolution data have been reported before. Our experimental results

seem to be inconsistent with the assumption that Sr would be distributed homogeneously

throughout the shell. Although Sr incorporation primarily depends on seawater concentration

and seawater concentration is conservative, some indication exists that physiological

processes may affect Sr incorporation. In a recent study, Elderfield et al. (2002) observed

decreasing Sr/Ca with increasing foraminiferal shell size and attributed this pattern to lower

calcification rates in larger shells. Similarly, Lea et al. (1999b) attributed incrasing Sr/Ca at

Publication III 65

higher temperatures and higher pH to higher calcite precipitation rates under these conditions.

Given that gametogenic calcite in G. sacculifer is deposited at a very high precipitation rate

(Bé, 1980), Sr/Ca may be enriched in outer calcite. Using this line of argument, the observed

increase in Sr/Ca upon dissolution may be due to a shift of the bulk shell chemistry to the

outer calcite. However, as long as we do not have Sr/Ca microprobe data for these shells, we

cannot generalize this hypothesis because other studies have shown that Sr/Ca may strongly

decrease in shells of crust-depositing foraminifera such as G. tumida (Brown and Elderfield,

1996) and O. universa (Haley and Klinkhammer, 2001). The different dissolution effects

observed in different foraminifera species are especially surprising with regard to the

relatively uniform Sr-concentrations in planktonic foraminifera (1.2-1.6 mmol/mol; e.g. Lea,

1999b). Although at this stage we cannot describe the dissolution behavior of Sr/Ca in various

foraminifera species from different locations by a single explanation, it is evident that Sr in

calcite is not necessarily less stable than Ca.

Martin et al. (1999) studied Sr/Ca over Quaternary glacial cycles and measured

variations in shells of N. pachyderma (sin.) and other planktonic foraminifera species that are

of the same order (~ +0.07 mmol/mol during glacials) as the subtle changes determined in our

study for dissolution. Dissolution effects thus can cause significant bias in reconstructions of

glacial/interglacial variations.

Stable isotopes

The variability in stable oxygen and carbon isotopes between samples is quite large.

However, the observed dissolution trends appear reasonable with regard to theoretical

considerations: As mentioned above, previous studies suggested that dissolution shifts bulk

shell δ18O and δ13C towards the isotope chemistry of the outer calcite (Lohmann, 1995;

Rosenthal et al., 2000). Due to the addition of gametogenic calcite in deeper water sections

where temperatures are colder, δ18O was observed to increase in foraminifera that were

subject to dissolution (Dekens et al., 2002; Erez, 1979b; Lohmann, 1995; Rosenthal et al.,

2000). Based on sediment material, Rosenthal et al. (2000) suggested that dissolution affects

δ18O and Mg/Ca similarly. In contrast, a recent study by Dekens et al. (2002) found that the

effect is not always the same for both proxies.

As expected from the lack of thermal and salinity gradients at the core site and the

constancy of the Mg/Ca data, G. sacculifer does not show a dissolution effect in δ18O.

However, the lower Mg/Ca in N. pachyderma (sin.) seems also to be reflected in increased

δ18O. The temperature underestimate due to dissolution changes in both parameters is similar

Publication III 66

in magnitude (~-2 K and -1.6 ± 1.2 K, respectively), but due to the large variability in δ18O it

is not possible to confirm or disprove any of the arguments suggested by Rosenthal et al.

(2000) and Dekens et al. (2002).

Available estimates on the impact of dissolution on δ13C in planktonic foraminifera are

based on the symbiont bearing G. sacculifer (Lohmann, 1995; Rosenthal et al., 2000) and the

symbiont-barren Globorotalia truncatulinoides (Lohmann, 1995). Both studies attributed

lowered δ13C values in the symbiont-bearing G. sacculifer to reduced symbiont

photosynthetic activity at depth. Our G. sacculifer data may indicate a subtle trend towards

lower values but given the large variability, only comparison with literature data allows to

conclude that δ13C of the inner calcite of symbiont-bearing foraminifera must be enriched so

that preferential dissolution of the inner calcite shifts δ13C to the lower values of the outer

calcite. However, in line with the results obtained by Lohmann (1995) on G. truncatulinoides,

δ13C in N. pachyderma increases upon dissolution. Lohmann (1995) already pointed to the

size related change in G. truncatulinoides' δ13C, with δ13C increasing in larger shells due to

some "vital effect". Recent laboratory culture experiments now allow to attribute part of the

ontogenetic isotope enrichment to seawater carbonate chemistry: Spero et al. (1997) and

Bijma et al. (1998) found δ13C and δ18O to increase in foraminiferal calcite if secreted under

lower [CO32-] (or pH). The effect was especially strong for the symbiont-barren G. bulloides

(Spero et al., 1997). Peeters et al. (2002) confirmed the laboratory experiments in a field study

and found that due to calcification at depth where [CO32-] (pH) is lower, δ13C in foraminiferal

shells may be higher although primary productivity and therefore the isotopic composition of

the seawater is lower. Similarly, the higher δ13C in stronger dissolved shells of N.

pachyderma (sin.) (and G. truncatulinoides) may be due to a shift of the bulk shell signal to

the chemistry of the outer calcite which is supposed to have been secreted under lower [CO32-

] (pH) deeper in the watercolumn (Kohfeld et al., 2000; Kohfeld et al., 1996). Unfortunately,

no carbonate chemistry data were found to calculate the [CO32-] (pH) over the calcification

depth range of N. pachyderma (sin.) so that we cannot quantify the theoretical considerations.

Considering the empirical δ44Ca-temperature equation by Nägler et al. (2000) for G.

sacculifer and similar results on the temperature dependency of δ44Ca in N. pachyderma (Zhu

and MacDougall, 1998), a lower isotopic composition would be predicted for partially

dissolved shells if the signal were shifted towards the chemistry of the outer calcite (reflecting

lower temperatures deeper in the watercolumn). Assuming that the temperature record would

Publication III 67

change similar to δ18O and Mg/Ca (-1.6 to -2 K) and that the δ44Ca - temperature slope for N.

pachyderma (sin.) is similar to the one determined for G. sacculifer (Gussone and Eisenhauer,

unpublished data), δ44Ca should be lowered by ~0.4-0.5‰ upon dissolution. In contrast, we

observe slightly increased values in N. pachyderma (sin.). Instead of selectively dissolving the

ontogenetic calcite, some process must lead to the selective removal of 40Ca. However, δ44Ca

appears constant in G. sacculifer. On the one hand, this result is in line with the constant

Mg/Ca and δ18O and seems consistent with the lack of thermal gradients at the core site. On

the other hand, if selective removal of the light isotope is active in N. pachyderma (sin.), the

number of 15 shells per analysis for G. sacculifer may be too small to yield a representative

average so that the actual effect could remain hidden behind the natural variability.

Whichever process is responsible, the observed variation in N. pachyderma (sin.) is within the

analytical uncertainty and negligible for the paleoceanographic record.

Because gametogenic calcite in G. sacculifer is added at depth where pH is lower,

lower δ11B values would be expected if dissolution shifted the signal to the chemistry of the

outer calcite. The observed increase in δ11B (from 20.8 to 22‰) is not consistent with the

expectation. Because the δ11B value of the most dissolved sample exceeds the estimates for

both, ontogenetic and gametogenic calcite (20.9 and 19.9‰, respectively; Fig. 20), selective

removal of the light isotope seems to occur similar to the observations on δ44Ca in N.

pachyderma (sin.). We also found reduced signal intensity during analysis of the dissolved

shells. This could be indicative for an overall leach-out of boron from the shell calcite during

dissolution. Although we do not have any boron concentration data to corroborate this

hypothesis, a previous study on boron preservation in oyster shells presented evidence for

preferential removal of lattice bound boron due to sub-aerial weathering (Cook, 1977).

Considering the microstructural breakdown presented in Figure 12, dissolution does not seem

to be restricted to the shell surface but an increasing number of fissures cut through the entire

shell and expose lattice areas in the interior of the shell wall. Due to the larger surface area

preferential removal of impurities is therefore possible even though they do not need to be

concentrated in particular inner or outer shell areas.

The observed change in δ11B (~+1.2‰ at 50% weight loss) is almost as large as the

glacial-interglacial (G-I) increase determined by Sanyal et al. (1995) for planktonic

foraminifera (1.8 ± 1‰). Although 50% weight loss would be a rather poor preservation for

sediment material, smaller dissolution mediated changes have the potential to affect δ11B.

Publication III 68

Conclusions

We analyzed the elemental and isotopic composition (Mg/Ca, Sr/Ca, δ18O, δ13C, δ11B,

δ44Ca) in partially dissolved shells of two species of planktonic foraminifera, G. sacculifer

and N. pachyderma (sin.) from well preserved sediments located in the Red Sea and the

Norwegian Sea, respectively. In summary, the chemical composition of the G. sacculifer

shells selected for this study reacts rather inert on dissolution. Except for Sr/Ca and δ11B no

significant chemical changes could be observed. While the lack of thermal and salinity

gradients at the core site could explain the constancy of temperature recorders such as δ18O,

Mg/Ca and δ44Ca, comparison with N. pachyderma (sin.) and previous dissolution studies

revealed that dissolution patterns are much more complex than so far assumed.

With regard to Mg/Ca, our calculated dissolution rates do not support the suggestion

that higher Mg concentrations significantly increase the dissolution susceptibility. No

preferential removal of Mg could be observed in G. sacculifer despite the fact that the element

is enriched in the outer calcite and overall more than 5 times more concentrated than in N.

pachyderma (sin.) which in turn shows a strong dissolution effect. Calculations of dissolution

rates lead us to conclude that the low Mg content in foraminiferal shells is not a major factor

controlling the overall dissolution behavior. Other processes have to be investigated instead.

For example, differences in the crystal structure and surface area in addition to variations in

the chemical shell composition could be responsible for inter- and intraspecifically variable

dissolution effects.

Although the Mg-content appears relatively insignificant to preferentially dissolve

Mg-rich calcite, Mg/Ca, δ18O and δ13C in N. pachyderma (sin.) are generally consistent with

a shift of the signal to the chemistry of the outer calcite. Supported by literature data

(Lohmann, 1995) we can conclude that δ13C of symbiont-barren species (N. pachyderma, G.

truncatulinoides) increases upon partial shell dissolution while δ13C of symbiont-bearing

foraminifera (G. sacculifer) decreases. With regard to δ18O, the temperature effect of colder

water masses predominates over carbonate chemistry and symbiont effects so that dissolution

increases δ18O independent of whether a foraminifera bears symbionts or not.

Given that the physico-chemical gradients at the Red Sea location are negligible and

that the measured Mg/Ca difference between inner and outer calcite of G. sacculifer is rather

small, our δ18O, δ13C and Mg/Ca results seem consistent with the hypothesis that ontogenetic

calcite is more prone to dissolution than gametogenic calcite (Brown and Elderfield, 1996;

Lohmann, 1995; Rosenthal et al., 2000; Russell et al., 1994). However, the dissolution

Publication III 69

mediated increase in δ11B and apparently also in δ44Ca is not consistent with a shift of the

signal to a deeper shell secretion environment. The observed change in δ44Ca is too small to

be significant for paleoreconstructions. In contrast, the increase in δ11B by ~1.2‰ is more

than half as large as the G-I increase determined by Sanyal et al. (1995) for planktonic

foraminifera. Although this implies a potentially strong dissolution-bias for δ11B, the number

of data presented herein is very small and needs to be confirmed in extensive sediment studies

and for other foraminifera species as well.

Sr/Ca was shown to increase in both species with progressing dissolution. The subtle

increase was found to be consistent with previous results on G. sacculifer (Bender et al.,

1975; Haley and Klinkhammer, 2002) but disagrees when compared to other foraminifera

species (Brown and Elderfield, 1996; Elderfield et al., 2000; Haley and Klinkhammer, 2002;

McCorkle et al., 1995). With the information currently available we have no explanation for

this result. However, with regard to N. pachyderma (sin.) the observed dissolution effect

could be significant for paleoreconstructions. All in all, the hypothesis that the most impure

calcite is more prone to dissolution (e.g. Brown and Elderfield, 1996) is not supported by our

Sr/Ca results.

Although we started off with the aim to quantify the impact of dissolution, effects

seem to vary among foraminifera species and the sites from where they were collected, i.e. the

physico-chemical gradients under which the shells were secreted. Consequently, the

prediction of chemical trends in progressively dissolving foraminifera shell assemblages is

not a straight-forward task. A single process - such as the assumption that ontogenetic calcite

dissolves faster than gametogenic calcite - is insufficient to explain the various dissolution

effects. Partial dissolution does not seem to be a simple shell thinning process but rather

increases the porosity of foraminiferal shells, exposing lattice areas which are protected in

undissolved shells. Increasing the shell porosity opens a pathway for the leach-out of certain

elements or isotopes in addition to the potentially higher solubility of inner shell calcite.

Individual dissolution patterns may also be complicated by trace metal interactions at crystal

surfaces, which may depend on relative concentrations in the calcite lattice and thus be

location-specific. In order to circumvent the complexity of post-depositional effects on

foraminiferal shell chemistry, examination of additional dissolution indicators is required to

identify the preservation state and minimize uncertainties. The microstructural breakdown of

G. sacculifer allows a first estimate of shell corrosion.

Publication III 70

Acknowledgements

Jan Helmke (GEOMAR, Kiel) and Jürgen Pätzold (GeoB, University of Bremen)

kindly provided foraminifera and core material. The experiments could not have been

successfully completed without the laboratory work of Georges Paradis and Dave Winter.

Pamela Martin shared her knowledge about trace element cleaning techniques and Michael

Kriews offered his laboratory facilities. We gratefully acknowledge laboratory help by Anja

Terbrüggen, Friedel Hinz, Carmen Hartmann, Silke Vetter and Folkmar Hauff. This research

was supported by NEBROC.

Publication IV 71

Publication IV

The impact of the ocean carbonate chemistry on living foraminiferal shell weight: A comment to Broecker and Clark’s „Carbonate ion concentration in glacial-age deep

waters of the Caribbean Sea“

Jelle Bijma, Bärbel Hönisch and Richard E. Zeebe

Geochemistry Geophysics Geosystems (in press, 2002)

.......................................................................................................................................................

Broecker and Clark (2002) use the “size normalized weight” of planktic foraminifera to

estimate the carbonate ion concentration ([CO3=]) of Atlantic glacial upper deep water. This

method was introduced by Lohmann (1995) and is based on the fact that, within a defined size

fraction, dissolution decreases shell weight in proportion to the degree of undersaturation with

respect to CaCO3. Based on the following assumptions,

1) the thickness of the foraminiferal shell wall does not depend on growth conditions,

2) the saturation [CO3=] for calcite, [CO3

=]sat., increases by 20 µmol kg-1 km-1 (Broecker

and Clark, 2001a),

3) the weight loss slope is universal and ca. 0.3 µg (µmol kg-1)-1 (Broecker and Clark,

2001a),

4) the offset between the bottom water and pore water [CO3=] was the same during

glacial time as during the Holocene,

Broecker and Clark (2002) calculate that the [CO3

=] of Atlantic glacial upper deep water was

14 µmol kg-1 higher than during the Holocene. Although they recognise that some of their

assumptions are not strictly valid, they do not assess the impact of those assumptions on their

[CO3=] estimate. Here we comment on several of those assumptions and attempt to quantify

their impact on Broecker and Clark's (2002) calculations.

Publication IV 72

Assumption 1

Broecker and Clark (2002) provide evidence that this assumption is not valid. At the

same pressure-normalised [CO3=], P. obliquiloculata from the Pacific Ocean is consistently

10 µg heavier than those from the Indian Ocean. (Spero and Lea, 1993) have shown that G.

sacculifer cultured under high light intensities grows bigger and is more massive than under

lower light conditions. Hemleben et al. (1987) have demonstrated that G. sacculifer cultured

at higher temperatures, grows larger. Unfortunately they did not measure shell weight.

Chamber number/size/weight relationships for the symbiont barren species G. bulloides differ

from location to location (Spero and Lea, 1996). For instance, comparable ontogenetic stages

from the Chatham Rise are bigger and heavier than those from the San Pedro Basin.

Apparently, besides potential genetic differences (e.g. Darling et al., 1999; Darling et al.,

2000; Huber et al., 1997), growth conditions affect the size normalised weight.

Broecker and Clark (2002) point out that temperature and [CO3=] are likely candidates.

Elderfield (personal communication) showed that shell wall thickness is closely related to

growth temperature. On the other hand, Bijma et al. (1999) demonstrated that shell weight of

O. universa is primarily a function of the [CO3=] of the ambient water (Figure 22a)2. At the

same [CO3=], O. universa shells grow heavier under high light than under low light

conditions. Apparently, the [CO3=] at the site of calcification (SOC), reflects the ambient

[CO3=] but is modified by physiological processes (respiration, calcification and

photosynthesis) of the foraminifer and its symbionts (Wolf-Gladrow et al., 1999a). Broecker

and Clark (2002) point out that in today’s ocean, a very tight correlation exists between

surface water [CO3=] and temperature and that core-top shell weights can thus not be used to

distinguish between a temperature or a carbonate ion dependence (in equilibrium with today’s

atmosphere, the temperature impact on [CO3=] varies roughly between 5 to 6 µmol kg-1 K-1

for surface alkalinities between 2100 to 2400 µmol kg-1, respectively).

2 It can be demonstrated that [CO3=] primarily affects shell thickness and that shell weight is a derived parameter of that relationship. For reasons of comparability and simplicity we use the [CO3=]-weight relationship here.

Publication IV

10

20

30

40

50

60

70

80

0 100 200 300 400 500 600 700

O. universa

shel

l wei

ght (

µg)

CO 32- (µmol kg -1 )

y = 29.5 + 0.051x R 2 = 0.67y = 23.19 + 0.11x R 2 = 0.55

73

0

20

40

60

80

100

0 100 200 300 400 500 600 700

G. sacculifer

shel

l wei

ght (

µg)


4 2

10

3

1

12

4 6 8

3

2

3 9

4 26

42

5

Figure 22. Planktonic foraminifera were grown in the laboratory in a 12hr light/dark cycle under a range of [CO3

=] until gametogenesis (for details see Bijma et al., 1999). Empty shells were dried, measured and weighed individually. A: Individual O. universa shell weights of the size fraction 500-600 µm plotted as a function of [CO3=]. The shaded area represents the range of ambient [CO3= ]. The data from several experiments (constant alkalinity, constant ΣCO2 and constant pH) are combined (Bijma et al., 1999). Linear regressions were fitted to data below and above ambient [CO3= ]. B: Globigerinoides sacculifer was only grown at constant ΣCO2. Average shell weights of three different size ranges (closed circles: 493-575 µm; open diamonds: 584-663 µm; closed squares: 762-845 µm). Numbers in the symbols indicate the number of specimens for each data point. Statistical analysis based on the individual data points shows that the slopes are significant at the 90%, 95% and 95% confidence level, respectively for the size intervals 493-575 µm, 584-663 µm and 762-845 µm. Regression analysis yields slopes of 2.5, 2.1 and 3.6 µg per 100 µmol kg-1, demonstrating that ontogeny has little effect on the relationship:

493-575 µm: Wgt (µg) = 27.3 + 0.025·[CO3=] (r2=0.39)

584-663 µm: Wgt (µg) = 41.4 + 0.021·[CO3=] (r2=0.22)

762-845 µm: Wgt (µg) = 59.8 + 0.036·[CO3=] (r2=0.28)

The sieve size range used by Broecker and Clark (350-415µm) can be converted to a real shell size range of 493-575µm by using the growth curve for G. sacculifer provided by Hemleben and Bijma (1994). In other words, the shell weight of G. sacculifer increases ca. 3µg for every 100µmol kg-1 increase in [CO3

=].

Publication IV 74

Barker and Elderfield (2002) have adopted an approach where they follow the evolution

of shell weight through time and compare this to predictions made from carbon system

modelling. They demonstrate that shell weight of G. bulloides decreases from the last

Termination towards the Holocene. This trend suggests that [CO3=] controls weight rather

than temperature since [CO3=] decreases while temperature increases during the deglaciation.

This finding is corroborated by laboratory experiments. We have evidence, albeit less well

constrained than for O. universa, that shell weight of G. sacculifer depends on the carbonate

chemistry of the ambient water (Figure 22b) and there is a priori no reason to assume that this

phenomenon is restricted to these two species or to foraminifera in general. A similar impact

of the carbonate chemistry has been demonstrated for corals (Gattuso and Buddemeier, 2000;

Gattuso et al., 1998; Kleypas et al., 1999) and for coccolithophorids (Riebesell et al., 2000;

Zondervan et al., 2001). The experimental results suggest that shell weight of individual G.

sacculifer increases by 3 µg for every 100 µmol kg-1 increase in [CO3=]. It should be noted

that only a part of the life cycle has been spent under controlled laboratory conditions and

hence that the real slope may be steeper. The point of the matter is that growth differences

have to be considered, not only between species and in space, but in time as well.

If we accept Sanyal et al.'s (1995) glacial surface water pH reconstruction, which is in

agreement with the ice-core pCO2 measurements and assume that the sites investigated by

Broecker and Clark (2002) were in equilibrium with the atmosphere, the glacial [CO3=] must

have been significantly higher. Depending on which scenario is followed to achieve glacial

pCO2 values, CO2 extraction or CaCO3 addition, the tropical glacial surface ocean [CO3=] was

50 to 120 µmol kg-1 higher, respectively, compared to the Holocene (Lea et al., 1999a).

Hence, before G. sacculifer settled to the ocean floor, glacial specimens must have been

heavier in weight than their Holocene counterparts from the same site. Using our empirical

relationship, 1.5 to 3.6 µg of the glacial weight increase in G. sacculifer was due to the fact

that shells grew heavier3. Hence, the effective glacial-Holocene weight difference decreases.

It is reasonable to assume that the change in the glacial surface [CO3=] was brought about by a

combination of both scenarios. Assuming that the average weight of glacial G. sacculifer was

2.6 µg heavier, the glacial interglacial weight difference for this species reduces from 4.9 to

2.3 µg. Consequently, using the 0.3 µg (µmol kg-1)-1 calibration of Broecker and Clark

(2002), the glacial [CO3=] increase reduces from 14 µmol kg-1 to 8 µmol kg-1.

3 Based on a glacial-interglacial temperature difference for tropical surface waters of 3 ºC, an average slope of [CO3

=] versus temperature of 5.5 µmol kg-1 K-1 and an experimental slope of 3µg weight increase for every 100µmol kg-1 increase in [CO3

=], the impact on shell weight of temperature alone is less than 0.5 µg.

Publication IV 75

Assumption 2 Although the true relationship between pressure and [CO3

=]sat. is exponential, a linear

approximation for the depth range between 3 to 4 km water depth is quite acceptable.

However, the coefficient of 20 µmol kg-1 km-1 (Ingle, 1975) used by Broecker and Clark

(2002) is the largest among a range of values. For instance, based on the relationship

[CO3=]sat. = 90·e(0.16·(z-4)) (Broecker and Takahashi, 1978), the change in the saturation [CO3

=]

between 3 to 4 km depth equals 13 µmol kg-1 km-1. Using the parameterisation of Millero

(1995), the saturation [CO3=] increases by 16 µmol kg-1 between 3 to 4 km depth. Although

Broecker and Clark (2002) acknowledge that the true slope may be in the range of 15 ± 2

µmol kg-1 km-1, they choose to use 20 µmol kg-1 km-1. Jansen et al. (2002) fitted the (more

convenient) equation of Broecker and Takahashi (1978) to the critical [CO3=] as calculated by

Millero (1995). Using this approximation ([CO3=]sat. = 88.7·e(0.189·(z-3.82))), the slope for the

depth range between 3 to 4 km is 16 µmol kg-1 km-1. A smaller pressure impact tends to

increase the weight loss slope and hence reduces the glacial [CO3=] estimate even more.

Assumption 3

Berger (1968) and Parker and Berger (1971) and many others after that, demonstrated

that planktic foraminifera are differently susceptible to dissolution. Hence, one could argue

that the critical [CO3=] is slightly different for each species and that therefore the weight loss

slopes must be species dependent and not universal. The fact that G. ruber does not show the

glacial-Holocene weight difference as the other species studied by Broecker and Clark (2002)

bolsters this contention. To verify the assumption that weight loss slopes are species

dependent we have replotted shell weights from Table 1 in Broecker and Clark (2001a)

against the pressure corrected [CO3=] (Figure 23). Based on the 20 µmol kg-1 km-1 change in

[CO3=]sat. used by Broecker and Clark (2001a), the slopes of the linear regressions are 0.46,

0.57 and 0.68 µg (µmol kg-1)-1 for G. sacculifer, P. obliquiloculata and N. dutertrei,

respectively. Using the smaller pressure effect on [CO3=]sat. of Jansen et al. (2002), the slopes

amend to 0.45, 0.66 and 0.65 µg (µmol kg-1)-1, respectively (Figure 23)44. Assuming an

average slope of 0.5 µg (µmol kg-1)-1, the estimated [CO3=] increase of glacial Atlantic upper

deep water, as calculated by Broecker and Clark (2001a), decreases by 43% from 14 to 8

4 Note that the inconsistency between the weight loss slopes and the ranking to dissolution provided by Parker and Berger Parker, F.L. and Berger, W.H., 1971. Faunal and solution patterns of planktonic foraminifera in

Publication IV 76

µmol kg-1. However, as argued above the slopes are species specific and ∆[CO3=] should

therefore be calculated on a per species basis. For the Caribbean cores the glacial-Holocene

increase in [CO3=] is then estimated to be 11 and 5 µmol kg-1 on the basis of G. sacculifer and

N. dutertrei, respectively. Apparently, the critical [CO3=] for G. sacculifer is higher than that

for N. dutertrei (i.e. at the same water depth, the Holocene-glacial weight difference for G.

sacculifer is larger than for N. dutertrei). The average of the two species is, of course, 8 µmol

kg-1 but the question arises which of the two species provides the best estimate?

Combining the impact of a steeper weight loss slope for G. sacculifer with that of

higher glacial surface water [CO3=] on G. sacculifer shell weight, reduces the glacial [CO3=]

increase estimate to 5 µmol kg-1.

Assumption 4

In addition to the assumption that the difference between the bottom and pore water

[CO3=] was the same during glacial time as during the Holocene, Broecker and Clark (2001a)

assume that the offset between bottom and pore water [CO3=] is constant with depth and

between different sites. However, because the rain ratio (carbonate carbon/organic carbon)

changes with depth and differs from location to location, these assumptions may not be valid.

In fact, Broecker and Clark (2001a) note that the observed correlation between the weights of

G. sacculifer, P. obliquiloculata and N. dutertrei for the core top samples from the Ceara Rise

demonstrates the variability of the [CO3=] offset between bottom and pore water.

Due to non-linear dissolution kinetics, the offset between bottom and pore water

[CO3=] changes drastically from above to below the saturation horizon (SH). Above the SH,

pore water is less saturated with respect to calcite than bottom water (due to respiration driven

pore water dissolution). Bottom water [CO3=] reaches the critical value at the SH; “interface”

dissolution starts and progresses exponentially towards greater depths. Because the

dissolution kinetics are not infinitely fast, an offset is created between the SH and the

lysocline. In this depth interval, called the transition zone, the saturation state of the pore

water increases from less saturated to more saturated than the bottom water5. This

demonstrates that the offset between bottom and pore water is not constant with depth and

surface sediments of the south Pacific. Deep-Sea Research, 18(1): 73-107. is most likely due to the combined effect of susceptability to dissolution and wall thickness (i.e. the initial shell weight). 5 Note that there is a depth where the [CO3

=] of bottom and pore water converge.

Publication IV

80

90

100

110

120

130

140

20 25 30 35 40 45

G.sacculifer

y = 35.51 + 2.22x R= 0.76

[CO

3= ] in s

itu n

orm

aliz

ed to

4 k

m d

epth

(µ m

ol/k

g)

Shell weight (µg)

77

70

80

90

100

110

120

130

140

20 25 30 35 40 45 50 55

P. obliquiloculata

y = 48.20 + 1.51x R= 0.71

[CO

3= ] in s

itu n

orm

aliz

ed to

4 k

m d

epth

(µ m

ol/k

g)

Shell weight (µg)

70

80

90

100

110

120

130

140

20 25 30 35 40 45 50

N. dutertrei

y = 50.32 + 1.54x R= 0.85

[CO

3= ] in s

itu n

orm

aliz

ed to

4 k

m d

epth

(µ m

ol/k

g)

Shell weight (µg)

Figure 23. Plots of whole shell weights (size range 355-415µm) as determined by Broecker and Clark (2002) versus the pressure corrected [CO3=], [CO3=]* = [CO3= ]in situ + ∆[CO3= ](4-z), where ∆[CO3= ](4-z) is the difference in [CO3=] between the water depth at the core site and 4 km and [CO3= ]at 4km depth is calculated after Millero (1995). Open circles represent Atlantic data, closed circles are from the Indian Ocean and + and x are from the western and eastern Pacific, respectively.

Publication IV 78

hence, that the slope of the weight-loss per unit change in [CO3=] changes below the SH.

Consequently, the size-normalised weight method should probably be restricted to cores that

have never seen in situ bottom water [CO3=] below the critical [CO3=].

One could argue that the variability in the [CO3

=] offset between bottom and pore

water is a fatal blow for Lohmann’s method. However, we should keep in mind that we are

dealing with a proxy and that such complications are to be expected. They basically set the

limit for the accuracy of the method. A better understanding of the [CO3=] variations between

bottom and pore water is needed to improve the robustness of the method.

Above the saturation horizon, the range of pH offsets between bottom and pore water,

for instance at the Ontong-Java Plateau, is somewhere between 0.02-0.04 pH units (Hales and

Emerson, 1996). This translates roughly to a [CO3=] offset between 5 to 10 µmol kg-1. Using

the average weight loss slope of 0.5 µg (µmol kg-1)-1, this implies that weight differences

between 2.5 to 5 µg are within the uncertainty for reconstructing bottom water [CO3=].

The restriction to cores that have never bathed in waters below the critical [CO3

=] has

also been noted by Broecker and Clark (2002). In addition, they argue that data from shallow

cores bathing in water with a [CO3=] higher than 120 µmol kg-1 should be omitted (because

the reduction in [CO3=] resulting from the release of respiration CO2 in the pore water is more

than compensated by the excess of bottom water [CO3=] over calcite saturation).

Some of the cores used by Broecker and Clark (2001a) to determine the weight loss

slopes are from below the SH (based on Jansen et al., 2002) or from in situ [CO3=] higher than

120 µmol kg-1. Limiting their data to [CO3=]sat. < [CO3

=]in situ < 120, the weight loss slopes for

G. sacculifer and N. dutertrei become 0.62 and 0.93 µg (µmol kg-1)-1, respectively. The

estimated glacial increase in [CO3=] for the Caribbean cores now decreases to 8 and 4 µmol

kg-1 on the basis of G. sacculifer and N. dutertrei, respectively.

If we combine the new estimate of the weight loss slope for G. sacculifer with the

impact of higher glacial surface water [CO3=] on this species, the ∆[CO3=] estimate based on

G. sacculifer reduces to 4 µmol kg-1. This brings the predicted average increase in [CO3=] of

Atlantic glacial upper deep water based on G. sacculifer (including the impact of higher

glacial surface water [CO3=] on initial shell weight) close to the prediction of 4 µmol kg-1

based on N. dutertrei (without an impact of higher glacial surface water [CO3=] on initial shell

weight). This could suggest that during growth the shell weight of N. dutertrei does not

Publication IV 79

respond strongly to the [CO3=], if at all. However, the predicted increase in [CO3=] of 4 µmol

kg-1 for Atlantic glacial upper deep water is clearly within the range of uncertainty related to

the variability in the [CO3=] offset between bottom and pore water leaving the question open

for further debate.

Acknowledgements

We would like to thank Howard J. Spero and David W. Lea for comments on earlier

versions of the manuscript and for many years of intensive collaboration in the laboratory in

which the shell weight measurements were generated.

Summary and outlook 80

3. Summary and outlook

To predict consequences of future climate variations especially with regard to

anthropogenic interference we have to understand the natural mechanisms that control the

climate system. The view into the past offers a rich source of information. However, the

quality of paleoceanographic reconstructions can be no better than the data base itself. This

thesis therefore deals primarily with the reliability of chemical proxies in planktonic

foraminifera shells as tracers of past seawater carbonate chemistry.

In the following section the main results are summarized and discussed in the light of

proxy applications in paleoceanography and paleoclimatology. Finally, perspectives for future

research are given.

3.1 Effects of symbiont photosynthesis and respiration on the stable boron isotopic

composition of foraminiferal shells

Culture experiments with living O. universa in combination with field-grown G.

bulloides and O. universa have shown that the pH in the microenvironment of planktonic

foraminifera is substantially affected by the photosynthetic sequestration of CO2, respectively

the release of CO2 during respiration processes. The observed difference in recorded pH

between shells grown under high light and low light conditions (∆pH ≈ 0.2) is only slightly

smaller than theoretical predictions based on microsensor pH-measurements (∆pH ≈ 0.34,

Rink et al., 1998) but significantly smaller than the model predictions cited in Publication I

(∆pH ≈ 0.4-0.76, Zeebe et al., subm.). New model runs have now shown that the difference

between model and culture data may be due to the 10-fold increased seawater boron

concentration used in the experiments (Zeebe et al., subm.). Taking into account that

increased boron concentrations modify the buffer capacity of the culture water, Zeebe et al.

(subm.) calculated different pH-gradients in the microenvironment compared to shells grown

in natural seawater. The smaller gradients calculated for shells grown in the modified

seawater are supposed to be responsible for the observed smaller differences between shells

grown under high light and low light. Upon correction of the boron concentration effect, the

model results are consistent with the culture data. This is a very promising result as it

demonstrates that the theoretical background for the use of δ11B as a paleoacidity-indicator is

well understood. The most important model prediction, however, is the constancy of the effect


over a wide seawater-pH range. This model outcome suggests that the reliability of δ11B is not

compromised by the physiological processes investigated in the culture experiments.

3.2 The effect of alkalinity on planktonic foraminiferal Ba/Ca

Theoretical arguments would have predicted higher Ba/Ca with increasing

precipitation rate. The arguments are based on experimentally determined similarities between

Sr2+ and Ba2+ incorporation into inorganic and biogenic calciumcarbonates (Lorens, 1981;

Morse and Bender, 1990; Tesoriero and Pankow, 1996; Zachara et al., 1991) and the finding

that Sr2+ uptake in planktonic foraminiferal calcite increases with pH, respectively

precipitation rate (Lea et al., 1999b). Similarly, Elderfield et al. (1996) developed a

biomineralization model for benthic foraminifera which argues for a calcification rate control

on trace element incorporation. However, the culture experiments carried out for Publication

II revealed slightly decreasing Ba2+ uptake in shells grown at higher alkalinities. Although the

observed effect is too small to significantly affect paleoceanographic reconstructions, it is

inconsistent with the theoretical predictions. A crucial point of this experiment, however, is

the use of planktonic and not benthic foraminifera which have so far been used for

reconstructions of glacial-interglacial variability (Lea, 1993; Lea and Boyle, 1990). Even

though planktonic foraminifera apparently react rather inert to changes in water column

alkalinity, benthic species may secrete their shells in environments which are actually

undersaturated with respect to calcite. It cannot be excluded that the incorporation of trace

elements under such conditions is different to the incorporation behavior of planktonic

foraminifera. For instance, (Marchitto Jr. et al., 2000) discovered that DZn and DCd in benthic

foraminifera may be affected by bottom water [CO32-] and argued that the supposed

dissolution effect on Ba/Ca in benthic foraminifera (Table 1, McCorkle et al., 1995) could be

equally due to an incorporation effect in undersaturated bottom water. In fact, preliminary

data on the benthic foraminifer Hoeglundia elegans indicate an anticorrelation between

bottom water alkalinity (or some associated parameter) and the Ba/Ca uptake ratio (G.J.

Reichart, unpublished data). This result is basically similar to the outcome of Publication II

and just as inconsistent with the theoretical predictions. However, it has to be pointed out that

H. elegans secretes an aragonitic shell and the result is thus not necessarily comparable to the

calcitic benthic foraminifera usually used for paleo-alkalinity estimates.

Laboratory experiments with benthic foraminifera are long-lasting and generally more

difficult than culture experiments with planktonic foraminifera. However, if additional data


prove the results described in Publication II and benthic coretop calibrations (G.J. Reichart,

unpublished data) to be true, alkalinity itself (or some associated parameter) may compromise

the use of benthic Ba/Ca as a tracer for past Ba2+ concentrations and paleo-alkalinity.

3.3 Changes in planktonic foraminiferal shell chemistry after partial dissolution in

undersaturated seawater

Previous research on the stability of chemical proxies in foraminifera is limited by a

number of uncertainties such as the ignorance of the actual dissolution state in sediments or

possible recrystallization processes. Simulated seafloor dissolution experiments carried out

for Publication III under controlled conditions in the laboratory and in combination with the

numerous proxies investigated allowed a detailed insight into the potential of partial shell

dissolution to change chemical proxies recorded in foraminiferal shells. Previous studies

focussed primarily on δ18O and Mg/Ca (Table 1). Although measured changes in the oxygen

isotopic composition of the shells studied herein were not significant, the observed variability

in δ18O and Mg/Ca is in line with known dissolution effects. However, the results presented in

Publication III should not be regarded as the final answer to all dissolution problems as

dissolution effects are very variable and have to be seen in the light of earlier research.

On the background of previous dissolution studies the results presented in Publication III

can be used to test different hypotheses previously put forward to explain observed effects.

Among these hypotheses the higher dissolution susceptibility of ontogenetic relative to

gametogenic calcite has the potential to explain the dissolution behavior of δ18O, δ13C, Mg/Ca

and maybe also Sr/Ca in shells of vertically migrating foraminifera that encounter physico-

chemical gradients during their ontogeny. However, δ44Ca and δ11B measured for this study

appear to increase upon dissolution and are therefore inconsistent with this model. δ44Ca and

δ11B have not been investigated in the light of partial dissolution before and additional studies

are required to confirm the observed trends. If dissolution shifts bulk shell chemistry to the

chemistry of calcite secreted at greater depth, δ44Ca and δ11B need to be significantly

dominated by other parameters which finally determine the relative stability of these elements

and isotopes in foraminiferal shell calcite.

Because it could be excluded that the higher Mg-content of certain shell parts may be

responsible for their lower stability, other parameters have to be considered to explain why

certain shell parts may be more soluble than others. For instance, gametogenic calcite is

precipitated much faster than ontogenetic calcite and different solubilities could be due to


differences in crystal structure or surface area. However, SEM examination of G. sacculifer

shells clearly showed that dissolution drives deep fissures into the shell surface,

demonstrating that dissolution also affects outer (gametogenic) calcite significantly and

exposes lattice areas to seawater which are protected in uncorroded shells. Based on these

results, partial dissolution does not seem to be a simple shell thinning process which would

argue for the outmost calcite layers to be removed first but increases the porosity, allowing

certain elements to be leached out of a shell.

The Sr/Ca data indicate that crystal impurities are not necessarily more prone to

dissolution. Although it was previously observed that Sr/Ca significantly decreases in

foraminifera species such as Orbulina universa (Haley and Klinkhammer, 2002), G. tumida

(Brown and Elderfield, 1996) and Globigerinoides ruber (D.W. Lea, personal

communication), other studies have shown subtle increases in Sr/Ca of G. sacculifer (Bender

et al., 1975; Haley and Klinkhammer, 2002). The different behavior found for different

foraminifera species can not be reconciled with the lack of pressure in laboratory dissolution

experiments and therefore has to be related to species-specific effects. Although higher

precipitation rates during gametogenic calcification could potentially explain why outer

calcite should be enriched in Sr and Sr/Ca therefore increases upon dissolution, comparison of

the herein observed dissolution pattern with previous studies remains puzzling. It could be

argued that in dependence of the overall elemental composition trace metals may interact at

mineral surfaces and thereby strengthen or weaken the stability of certain elements in the

crystal lattice. Considering the extent of vertical migrations, the encountered physico-

chemical vertical gradients and the overall specific seawater chemistry at different study sites,

dissolution effects may not only be species- but also location-specific (see also Table 1). In

summary, it has to be concluded that the process of partial dissolution is not yet very well

understood and that the prediction of dissolution effects is anything but a straight-forward

task. Without a fundamental understanding of the respective dissolution processes it is not yet

possible to quantify the extent to which certain proxies change upon dissolution and to correct

proxies measured in apparently corroded shells.

3.4 Foraminifera collected from sediment cores - identifying their preservation state Publication III has demonstrated that corrosion of planktonic foraminifera is reflected

in their shell microstructure and can easily be detected by SEM examination. This is an

important finding as it opens a simple way to evaluate the preservation state of shells used for

chemical analyses or bottom water acidity estimates. As discussed in Publication IV,


planktonic foraminiferal shell thickness or density is not only determined by post-depositional

dissolution but also varies with changing calcite saturation in surface waters. As demonstrated

herein, visual inspection of the shells allows to differentiate between real corrosion and less

favorable growth conditions. Comparable microstructural dissolution patterns have already

been observed by Volbers and Henrich (2002) for G. bulloides and Helmke and Bauch (2002)

for N. pachyderma (sin.). Compared to labor-intensive estimates of weight loss or break-up of

shells, SEM examination can be done on single shells and thus offers a relatively fast method

to judge whether the primary chemical signal of numerous downcore samples may have been

altered post-depositionally or not.

3.5 Perspectives for future research

The present study investigated the chemical composition of foraminifera shells in

response to changing seawater carbonate chemistry. Some open questions were answered,

however, as it is always in investigations of this kind, a wide field of new questions has been

opened as well. The boron isotopic composition of cultured and field-grown foraminifera

presented in Publication I was found to be significantly offset to previously established

empirical calibration curves. The reason for the observed offsets between studies and species

is most probably due to differences in the analytical methodology. Although it was assumed

that the differences between samples of the same species that are measured in the same

laboratory reflect the conditions under which the foraminifera were grown, differences and

uncertainties between laboratories and species are unsatisfactory with regard to

paleoceanographic reconstructions. The systematic reassessment of the analytical

methodology is therefore planned in the near future. In addition, we intend to produce and

establish an analytical standard with a boron isotopic composition and carbonate matrix

similar to the microfossils studied for paleoceanographic reconstructions. Foraminifera will

also be investigated in the light of changing salinity and temperature, parameters that have

recently been suggested to affect the boron isotopic fractionation during foraminiferal shell

secretion. This work will not be restricted to planktonic foraminifera alone, as there is a

pronounced need to investigate benthic foraminifera as well. Empirical calibration curves for

single benthic species are required which should be established preferentially in laboratory

culture experiments.

Publication II and III have shown that the mechanisms operating during trace element

incorporation and selective removal upon partial dissolution are still poorly understood.

Modeling studies in combination with microscale measurements of single foraminifera shell


walls are planned or partly underway in order to better understand the data obtained in this

and other studies. Cooperations with mineralogists and chemists are intended in order to

determine shell surface areas, yield a better understanding of how crystal structure could

affect the calcite stability and to find out which elements have the potential to interact and

thereby vary dissolution susceptibility.

The dissolution effect on δ11B clearly needs to be confirmed. As boron in calcite

occurs in trigonal and tetrahedral coordination states (Sen et al., 1994), and the isotopic

composition of each coordination state may be analog to the isotopic composition of trigonal

and tetrahedral boron in solution (personal communication N.G. Hemming), preferential

removal of one of the coordinations could possibly explain a change in isotopic composition

upon dissolution. Depending on which coordination is less stable, such a process would

remove isotopically light boron (if tetrahedral coordination is less stable) and leave the

residual with a heavier δ11B or vice versa (if trigonal coordination is less stable). Following

this assumption and given that boron in aragonite (corals) only occurs in the tetrahedral

coordination (Sen et al., 1994), no dissolution effect would be expected for aragonite. This

prediction could be easily tested in an experiment. However, this consideration clearly

depends on the assumption that only B(OH)4- (and no B(OH)3) adsorbed at the surface of a

foraminifera shell gets incorporated into the lattice (otherwise corals and foraminifera would

have a very different isotopic composition than shown e.g. in Hemming and Hanson, 1992)

and that a fractionation accompanies the structural transformation which isotopically enriches

the trigonal coordination state relative to the tetrahedral without changing the overall shell

isotope value (Sen et al., 1994). Because boron is assumed to substitute at the anion site in

CaCO3 and the CO32--site in calcite is larger than in aragonite (Sen et al., 1994), the

occurrence of the smaller trigonal coordination state in calcite could indicate that it is better

embedded in the calcite lattice than the larger tetrahedral coordination. Using this line of

argumentation, the observed increase in δ11B (Publication III) could be due to the higher

stability of the isotopically heavier trigonal coordination state in foraminiferal calcite.

However, this process would have to dominate over the preferential removal of inner

compared to outer calcite (as suggested for δ18O, δ13C and Mg/Ca) and the expected lower

δ11B incorporated at depth. As a first start, shells of Globigerinella equilateralis, picked and

partially dissolved in addition to the G. sacculifer shells from the Gulf of Aqaba core site

(Publication III), are available to challenge the G. sacculifer results. Furthermore, sediment

studies are planned with shells from different water depths.


Finally, we are also interested in the expansion of our knowledge to other marine

calcifiers. A unique set of hermatypic corals cultured under variable conditions of seawater

carbonate chemistry has already been analyzed for their boron isotopic composition. The

response to pH is significant, although different to the fractionation observed in foraminifera.

To complete this data set, the samples are already in preparation for minor and trace element

analyses such as Sr/Ca, Mg/Ca, U/Ca and Ba/Ca. Although the incorporation of these

elements into coral aragonite has not yet been calibrated with respect to changes in seawater

carbonate chemistry, evidence already exists that the chemical composition of coral skeletons

may respond in a manner similar to planktonic foraminifera (e.g. Cardinal et al., 2001; Cohen

et al., 2002; Kühl et al., 1995; Min et al., 1995; Shen and Dunbar, 1995).

Appendix 87

4. Appendix

Reconstructing and modeling past ocean carbonate chemistry –

Working Group 3 report of the ESF Explanatory Workshop on "The ocean carbon

cycle and climate change", Delmenhorst, September 1-4, 2001

Bärbel Hönisch (rapporteur), Jonathan Erez (discussion leader), Christina Crone, Heiko

Jansen, George P. Lohmann, Guy Munhoven, Martin R. Palmer, Ann D. Russell, Dieter A.

Wolf-Gladrow, Richard E. Zeebe, Patrizia Ziveri

.......................................................................................................................................................

Abstract

The chemical composition and preservation state of biogenic calcium carbonates

derived from marine sediments are routinely used for paleoceanographic and

paleoclimatologic reconstructions despite the fact that their calibration to environmental

parameters is often a matter of debate. Significant differences are observed between

laboratory culture experiments, field calibrations (as obtained from different locations and

depths in the water column) and core top-based analyses. The reasons for these discrepancies

are not well understood and call for further evaluation. Working group 3 reconsidered the

proxies currently available for reconstructions of the oceanic carbonate system. In addition to

the prevalent recorders of past ocean chemistry, foraminifera, we also discussed

coccolithophorids. Previous carbonate system reconstructions and limitations of several

proxies are summarized and suggestions for future research are proposed.

Appendix 88

Introduction

On glacial-interglacial time scales oceanic carbonate chemistry determines

atmospheric pCO2. However, the underlying mechanisms and possible feedbacks with climate

change are still not well understood. Knowledge of the nature and amplitude of natural

fluctuations in the past can be used to assess the stability of modern subsystems and their

potential range of variations in the future. Understanding the climate system therefore

requires the reconstruction of physical, chemical and biological parameters that characterize

the ocean carbonate system over glacial and interglacial time scales as well as the transitions

between them. Over the past decade a number of proxy relationships based predominantly on

foraminifera have been established on the basis of laboratory and field experiments. Among

others, the stable boron isotopic composition in foraminiferal shells was found to record

marine pH (Spivack et al., 1993), Ba/Ca was used to infer alkalinity (Lea and Boyle, 1989)

and differences between the influence of [CO32-] on the stable oxygen and carbon isotopic

composition of Globigerinoides sacculifer and G. ruber was found to be useful for past

carbonate ion concentration estimates (Bijma et al., 1999; Spero et al., 1999). More recently,

some new proxies such as foraminiferal U/Ca (Russell, 2001), S/Ca (Erez et al., 2001) and the

CaCO3 size fraction index (Broecker and Clark, 1999) have been found to record [CO32-].

In addition to these chemical proxies of the seawater carbonate system, the

preservation state of carbonates can be used to estimate bottom water undersaturation for

aragonite and calcite. One approach is to examine the depth of their lysoclines, which can be

defined as the levels of maximum solution rate increase in the deep sea (Berger, 1968).

Relating the preservation state of carbonates in the sediment to the saturation state of bottom

water yields information on [CO32-]in situ. For example, the dissolution-driven shell-thinning of

planktonic foraminifera was found to give a good approximation of bottom water [CO32-]

(Broecker and Clark, 2001a; Lohmann et al., 1999).

Unfortunately, the interpretation of these proxies is often hampered by their

dependency on additional variables. Species specificity, vital effects of the organisms, and

even the susceptibility to diagenesis complicate data interpretation. Hence, proxy

relationships are not as simple as we would like them to be. For example, the interpretation of

geochemical proxies in surface-dwelling planktonic foraminifera is complicated by the

presence of a significant fraction of calcite added at depth. Thus, since most calibrations are

empirical, a certain discrepancy exists between laboratory experiments and the real situation

at the seafloor. Combination of these proxy development approaches with better

Appendix 89

understanding of the calcification mechanisms and numerical models should yield higher

reliability of the proxies. Eventually, such new data and its proper modeling would render

further insight into the role and the impact of the carbon cycle on climate oscillations and in

particular resolve the mechanisms that control the operation of the oceanic carbon cycle.

Working group 3 discussed the carbonate system proxies (Table A1) and their

limitations. In order to obtain high-quality paleoreconstructions, future research needs to

focus on the removal of these uncertainties. Recommendations point towards field

investigations and culture experiments, as well as towards numerical models and the

improvement and further development of existing and new analytical techniques.

2. Proxies for ocean carbonate chemistry and their limitations 2.1 Carbonate chemistry 2.1.1 Ba/Ca to infer alkalinity

Ba is a nutrient-like tracer similar to Cd and δ13C, because biological activity extracts

these elements from surface waters and gravitation transfers them toward the seafloor in

sinking particles. On its way from the North Atlantic to the deep North Pacific, deep water is

progressively enriched in Ba. The close correlation between Ba and alkalinity in seawater

(Chan et al., 1977; Lea and Boyle, 1989) is mechanistically not well understood (Bishop,

1988; Chan et al., 1977; Chow and Goldberg, 1960; Lea, 1993), but is thought to be related to

the simultaneous release of alkalinity through CaCO3 dissolution and regeneration of Ba at

the seafloor. However, Lea (1993) suggested that changes in the thermohaline circulation

redistribute Ba and alkalinity similarly, thereby allowing reconstruction of past alkalinity

distributions from benthic foraminiferal Ba/Ca.

The main limitation of Ba as a paleoproxy is due to the short oceanic residence time

on the order of 10,000 years (Broecker and Peng, 1982; Chan et al., 1977). The Ba-alkalinity

correlation is not perfectly applicable on a time scale longer than this period. However, the

fact that Ba is incorporated into foraminifera shells in direct proportion to the seawater

concentration (Lea and Boyle, 1989; Lea and Spero, 1992; Lea and Spero, 1994) allows us to

estimate paleo-Ba concentrations from foraminifera deposited in sediments. Independent

estimates of seawater carbonate chemistry would offer an opportunity to verify whether the

present-day slope of the Ba-alkalinity relationship is applicable to the past as well. A

multiproxy approach would provide the best means of calculating alkalinity for various time

Tab

le A

1. R

econ

stru

ctin

g pa

st o

cean

car

bona

te c

hem

istry

: pro

xies

, lim

itatio

ns a

nd e

stim

ates

. Ple

ase

note

that

the

tabl

e m

erel

y pr

esen

ts p

roxi

es th

at

wer

e di

scus

sed

durin

g th

e w

orks

hop.

Pr

oxy

targ

et

para

met

er

inte

rfer

ing

para

met

ers

estim

ate

for

LG

M

limita

tions

re

fere

nces

Ba/

Ca

se

awat

er[B

a2+] a

nd

alka

linity

none

on

plan

kton

ics;

pr

essu

re o

n be

nthi

c D

Ba

+20-

25 ±

10

µmol

kg-1

(C

ircum

pola

r Dee

p W

ater

)

shor

t res

iden

ce ti

me,

m

echa

nist

ic li

nkag

e be

twee

n [B

a2+] a

nd a

lkal

inity

not

un

ders

tood

Lea

and

Boy

le 1

989;

Lea

199

3;

Lea

and

Sper

o 19

92, 1

994

δ11B

pH

spec

ies-

spec

ific,

sym

bion

ts, m

aybe

te

mpe

ratu

re a

nd e

arly

di

agen

esis

+0.2

± 0

.1 u

nits

(s

urfa

ce w

ater

) +0

.3 ±

0.1

uni

ts (d

eep

ocea

n)

mat

eria

l and

tim

e co

nsum

ing,

onl

y m

ixed

be

nthi

cs

Sany

al e

t al.

1995

, 199

6, 2

001

Car

bona

te io

n ef

fect

on

δ13C

[C

O32-

] de

pth

habi

tat (

G.

sacc

ulife

r and

G.

rube

r)

+55

± 63

µm

ol k

g-1

(Ind

ian

Oce

an)

Sp

ero

et a

l. 19

97,1

999;

Bijm

a et

al.

1999

; Rus

sell

and

Sper

o 20

00

U/C

a [C

O32-

] sp

ecie

s-sp

ecifi

c,

tem

pera

ture

? +8

0-10

0 ±

60

µmol

kg

-1

Mn

cont

amin

atio

n R

usse

ll et

al.

1994

; Rus

sell,

200

1

SO42-

/CO

32-

[CO

32-]

Mg,

tem

pera

ture

eq

uiva

lent

to +

0.2

pH

units

(dee

p oc

ean

) B

aSO

4 con

tam

inat

ion

Erez

et a

l., 2

001

Size

-no

rmal

ized

w

eigh

t

deep

[CO

32-]

diss

olut

ion

offs

et b

etw

een

pore

w

ater

and

bot

tom

w

ater

satu

ratio

n,

grow

th c

ondi

tions

no d

ata

loca

l cal

ibra

tions

requ

ired,

re

prod

ucib

ility

of w

eigh

t m

easu

rem

ents

: ± 6

%

Lohm

ann

1995

; Loh

man

n et

al.

1999

; Ros

enth

al e

t al.

2000

; B

roec

ker e

t al.

2001

; Bro

ecke

r and

C

lark

200

1 si

ze fr

actio

n in

dex

deep

[CO

32-]

diss

olut

ion

no

dat

a

(± 5

µm

ol k

g-1)

non

calc

ite m

ater

ial t

rapp

ed

in sh

ells

, Cor

g rai

n,

cont

inuo

us b

reak

up a

fter

buria

l?

Bro

ecke

r and

Cla

rk, 1

999

R

efle

cten

ce/

light

ness

+

wei

ght

corr

osio

n

qu

alita

tive

estim

ates

only

Hel

mke

and

Bau

ch, 2

002

Appendix 91

scales which can then be compared with the Ba concentrations recorded in foraminiferal

shells.

2.1.2 δ11B to infer pH

Inferring seawater pH from stable boron isotopic compositions is based on the isotopic

fractionation between dissolved boron in seawater and boron in CaCO3. The uncharged boron

species B(OH)3 is enriched in 11B by ~20‰ over the charged species B(OH)4-. Of these two

dominant aqueous species, it is the charged borate that is incorporated into carbonate

minerals. As the fraction of B(OH)4- and B(OH)3 changes with pH, so must their respective

isotopic compositions. The isotopic composition of boron in CaCO3 therefore is enriched with

the heavier isotope 11B with increasing pH (Hemming and Hanson, 1992; Sanyal et al., 2000).

The boron isotopic composition in carbonates is also highly sensitive to local variations in

pH. The calcification process itself and microenvironments (i.e. associations with symbionts

15

20

25

30

35

40

7.5 8 8.5 9 9.5

inorg. precipitation

cultured O. universa

B(OH)4-

Holocene G. sacculifer

Holocene O. universa

cultured G. sacculifer

δ11B

(‰)

pH

modern oceanic range

T = 25°C & S = 35‰(pK = 8.60; α = 19.4‰)

Figure A1. Present state of the δ11B proxy calibration. Red: theoretical curve according to Kakihana et al. (1977); black: inorganic precipitation results (Sanyal et al., 2000); blue: Orbulina universa from culture experiments (closed circles, Sanyal et al., 1996) and core-top sediment samples (open circles, Sanyal et al., 1997); green: Globigerinoides sacculifer from culture experiments (closed triangles, Sanyal et al., 2001) and core-top sediment samples (open triangles, Sanyal et al., 1995).

Appendix 92

or precipitation within extrapallial fluids) must therefore be considered. While the planktonic

foraminifer G. sacculifer and benthic foraminifera appear to incorporate δ11B with little or no

fractionation compared to the theoretical curve, O. universa shows an offset from those

foraminifera by ~3.3 ‰ (Sanyal et al., 2001). The offset was suggested to be due to a vital

effect, although its nature could not be explained: both planktonic species are spinose and

symbiont-bearing and should therefore react similarly.

The effect of symbiont photosynthesis has recently been investigated in a diffusion-

reaction model by Zeebe at al. (subm.). They calculated a constant but significant offset

between δ11B in planktonic foraminiferal calcite and the isotopic signature of B(OH)4- in the

seawater medium. Recent laboratory culture data (Hönisch et al., subm.) are in good

agreement with the model results.

Stable boron isotopic analyses, using negative thermal ionisation mass spectrometry

(nTIMS), have several complications. First of all, the technique requires several hours of

permanent operator assistance and numerous replicate analyses until accurate values can be

obtained. Second, to achieve a reproducible result, approximately 4 ng boron are required per

analysis. Since foraminifera contain 5-15 ppm B (Hemming et al., 1998), up to 10 mg

foraminiferal calcite are needed per sample (when considering weight loss during cleaning

and multiple replicate analyses). Especially for the investigation of deep water chemistry the

second point is crucial, as the abundance of benthic foraminifera is too low to routinely allow

single-species analyses. Sanyal et al. (1995) therefore combined several species for their

deepwater record despite possible differences in habitat (epifaunal/infaunal) characterized by

a range of pH conditions, and species-specific offsets like the ones found for planktonic

foraminifera (Sanyal et al., 2001). These factors (Sanyal et al., 1997; Sanyal et al., 1996) may

have biased the obtained value, which suggested a 0.3 pH units increase for last glacial

deepwater (Sanyal et al., 1995).

In order to solve the problems named above, it is desirable to reduce the amount of

material required for analyses, to speed up measurements and to generally expedite the

analytical procedure so that the investigation of past ocean acidity can be realised extensively

in the future.

Appendix 93

2.1.3 Deconvolution of the carbonate ion effect to infer [CO32-]

This approach is based on the deconvolution of foraminiferal δ13C records to calculate

the change in surface [CO32-] and δ13CΣCO2 through time: The stable carbon and oxygen

isotopic compositions of planktonic foraminifera decrease with increasing carbonate ion

concentration (Bijma et al., 1999; Spero et al., 1999). Among the investigated planktonic

foraminifera, G. sacculifer and G. ruber share the same habitat but the slope in δ13C vs.

[CO32-] is twice as large in G. ruber as in G. sacculifer. This species-specific difference is

used to distinguish between the effect of [CO32-] and a simultaneous change in δ13CΣCO2.

Application to the sediment record leads to the estimate of +55 ± 63 µmol kg-1 [CO32-] for the

Indian Ocean during the last glacial (Spero et al., 1999). Unfortunately, this method is

restricted to tropical surface waters, where G. sacculifer and G. ruber occur.

2.1.4 U/Ca to infer [CO3

2-]

Laboratory experiments revealed that U/Ca in planktonic foraminifera shells is

inversely related to [CO32-] (Russell, 2001). The symbiont-barren G. bulloides incorporates

approximately twice U/Ca than the symbiont-bearing O. universa at the same [CO32-]. No

consistent temperature effect on the record has been found above 19°C. Application of the

U/Ca relationship to Caribbean cores suggested that glacial [CO32-] was 80-100 ± 60 µmol kg-

1 higher than during the Holocene.

Although the approach is generally promising, the study of several sediment cores

revealed that contamination by Mn-carbonates places a significant diagenetic overprint on the

incorporated U/Ca which may limit the general applicability of this proxy to sediments above

the redox front.

2.1.5 SO4

2-/ CO32- to infer [CO3

2-]

In laboratory culture experiments, Erez et al. (2001) observed a constant distribution

coefficient between SO42-/CO3

2- in the shells of benthic and planktonic foraminifera and SO42-

/CO32- ratio in seawater. Since the seawater SO4

2- inventory is not expected to have changed

on glacial-interglacial time scales, [CO32-] can be reconstructed. In situ calibrations of this

proxy in the Gulf of Eilat gave similar results to those of the laboratory experiments.

However, in the Little Bahama Bank a temperature effect was revealed which may have been

caused by changes in the Mg ion content, apparently affecting the SO42- content of

foraminiferal shells. Correction of this temperature effect leads to the empirical negative

Appendix 94

correlation between seawater [CO32-] and foraminiferal SO4

2-/CO32- as found in laboratory

culture experiments. Preliminary comparisons of SO42-/CO3

2- from Holocene and glacial

benthic foraminifera show variability in pH similar in magnitude to that estimated

independently from δ11B (approximately 0.2 pH units increase in the glacial deep Pacific).

The advantage of this proxy is the very small sample size required for routine

measurements using a Magnetic Sector ICP-MS. Hence it is practical to be used for benthic

foraminifera from deep sea sediments. However, the proxy is still under development and not

much is known about its limitations. One possible interference may be the contamination with

extraneous phases like barite (BaSO4).

2.2 Carbonate preservation 2.2.1 Size-normalized shell weight

The average mass of planktonic foraminifera is primarily determined by their size, but

there is a measurable secondary relationship of shell mass to water depth (Lohmann, 1995;

Lohmann et al., 1999; Rosenthal et al., 2000). Due to dissolution, the size-normalized mass of

nearly all species is lower in deeper water than it is in shallow water, and the decrease is

continuous over a wide range of carbonate saturation states, even well above the calcite

lysocline. Based on shells of the three species G. sacculifer, Pulleniatina obliquiloculata and

Neogloboquadrina dutertrei, Broecker and Clark (2001a) determined an average weight-loss

slope of 0.3 ± 0.05 µg (µmol kg-1)-1 decrease in pressure-corrected deep sea carbonate ion

concentration. This relationship allows estimates of changes in seawater carbonate content

from the size-normalized mass of planktonic foraminifera.

To use the relationship as a paleocarbonate ion proxy, this method requires that the

offset between pore and bottom water saturation is constant. However, numerous

investigations (Archer et al., 1989; Berelson et al., 1990; Berelson et al., 1994; Hales and

Emerson, 1996; Hales and Emerson, 1997a; Jahnke et al., 1994; Jahnke et al., 1997) have

applied microelectrodes and benthic flux chambers to validate the theory of respiration-driven

dissolution in-situ (Table A2, see also section 2.3.1). They conclude that 40-60 % of the

calcite dissolution above the saturation horizon can be attributed to metabolic processes. The

amount of organic matter reaching the seafloor varies between sites and depends on depth.

Assuming increased productivity on glacial time scales, the magnitude of this effect might

have been even stronger. Application of this proxy should therefore be restricted to locations

where strong changes in paleoproductivity are not expected.

Tab

le A

2. I

n si

tu in

vest

igat

ion

of se

dim

enta

ry c

arbo

nate

dis

solu

tion.

R

efer

ence

L

ocat

ion

Dep

th (m

)

∆ (µ

mol

l-1) o

r Ω

δ r

M

D1 (

%)

MD

2 (%

)

Hal

es e

t al.

(199

4)

Nor

th A

tlant

ic

2100

-540

0

Ω =

-0.2

– +

1.75

35 –

67

> 60

Ja

hnke

et a

l. (1

994)

Ea

ster

n N

orth

Atla

ntic

31

00

∆ =

+27

0

noda

taA

rche

r et a

l. (1

989)

Eq

uato

rial A

tlant

ic

3800

-500

0

∆ =

-27

– +9

so

me

som

eH

ales

and

Em

erso

n (1

997)

W

este

rn e

q. A

tlant

ic (C

eara

Ris

e)

3300

-470

0

Ω =

+0.

8 –

+1.2

> 20

> 45

M

artin

and

Say

les (

1996

) W

este

rn e

q. A

tlant

ic (C

eara

Ris

e)

3200

-470

0

∆ =

-28

– +1

3 0.

8 –

1.4

36 –

66

37 –

92

Ber

elso

n e

t al.

(199

0)

Cen

tral e

q. P

acifi

c 44

00-5

000

∆

= -6

– +

12a

no

dat

aso

me

B

erel

son

et a

l. (1

994)

C

entra

l eq.

Pac

ific

3380

-456

0

Ω =

+0.

78 –

+0.

87

no

dat

a60

– 1

00

Cai

et a

l. (1

995)

C

alifo

rnia

con

tinen

tal m

argi

n 41

00

Ω =

+0.

69

1.0

no d

ata

som

e

Jahn

ke e

t al.

(199

7)

Cal

iforn

ia c

ontin

enta

l mar

gin

800-

3700

∆

= -1

8 –

-9

3.6

– 7.

3 no

dat

a90

H

ales

and

Em

erso

n (1

996)

W

este

rn e

q. P

acifi

c (O

nton

g Ja

va P

late

au)

2300

-300

0

Ω =

+0.

75 –

+0.

91

no

dat

a>

65

Jahn

ke e

t al.

(199

4)

Wes

tern

eq.

Pac

ific

(Ont

ong

Java

Pla

teau

)30

00, 4

400

∆

= -6

, -37

no d

ata

som

e

∆ =

[CO

32-] in

situ

- [

CO

32-] c

mea

sure

s th

e sa

tura

tion

stat

e, w

here

[C

O32-

] c . A

ltern

ativ

ely,

Ω =

[C

O32-

] in s

itu -

[C

O32-

] c . The

rai

n ra

tio C

org/C

aCO

3 is

de

note

d by

δr.

MD

1 de

note

s th

e fr

actio

n of

car

bona

te a

bove

the

satu

ratio

n ho

rizon

that

is d

isso

lved

. MD

2 de

note

s th

e co

ntrib

utio

n of

met

abol

ic

proc

esse

s to

carb

onat

e di

ssol

utio

n be

low

the

satu

ratio

n ho

rizon

. a H. J

anse

ns’s

cal

cula

tion

usin

g G

EOSE

CS

data

, the

val

ues a

re n

ot m

entio

ned

in th

e pa

per.

Appendix

10

20

30

40

50

60

0 100 200 300 400 500 600 700

G. sacculifersh

ell w

eigh

t (µg

)


(6)

(4)

(5)

(2)

(4)

(2)

y = 27.3 + 0.025x R 2 = 0.39

96

10

20

30

40

50

60

70

80

0 100 200 300 400 500 600 700

O. universa

shel

l wei

ght (

µg)


y = 29.5 + 0.051x R 2 = 0.67y = 23.19 + 0.11x R 2 = 0.55

Figure A2. Increased foraminiferal shell weight under higher [CO32-] during shell growth: a)

Linear fit regression for G. sacculifer at 29°C, real size range: 493-575 µm. Numbers in brackets represent the number of shells per average. Data compiled from several laboratory culture experiments. b) Linear fit regression for O. universa at 22°C, real size range: 500-600 µm. The shaded area represents the range of ambient [CO3

2-]. Data modified after Bijma et al. (1999).

Complication also arises from observations on an initial increase in the average shell

weight of a freshly sedimented foraminiferal population: At the initial stages of dissolution

(probably still above the lysocline) the thinshelled, light-weighted individuals disintegrate

first. The loss of these shells from the sedimentary community is consistent with an observed

increase in foraminiferal δ18O and δ13C which can be attributed to the lighter isotopic

composition of thinshelled individuals (Erez, 1979a; Erez, 1979b). Lohmann‘s data indeed

show the maximum weight per individual slightly above the lysocline. It should also be noted

that the dissolution rate of foraminifera is not a linear process. This is due to the exponential

increase of the specific surface area of a foraminifera with dissolution (Honjo and Erez, 1978)

and to the non-linear nature of calcite-dissolution as a function of undersaturation (Keir,

1980).

An intriguing factor is also the dependency of the proxy on the target parameter itself,

i.e. deep [CO32-] does not exclusively determine the preservation state of foraminiferal shells

in the sediment but shell weight also depends on surface [CO32-] experienced during shell

precipitation. Figure A2 demonstrates this dependency for the species G. sacculifer and O.

universa. For a glacial increase in surface water [CO32-] on the order of 55-100 µmol kg-1 (as

determined from U/Ca and carbonate ion effect, Table A1), the shell weight of G. sacculifer

Appendix 97

thus increases by approximately 1.6–2.9 µg – independent of concomitant changes in deep

water saturation. A detailed examination of available culture and sediment trap data is

required to better estimate the magnitude of the observed growth effects and to determine how

temperature affects the carbonate dependent growth variability at a certain locality over

glacial/interglacial time scales.

2.2.2 Reflectance/ lightness of foraminiferal shells

A qualitative estimate of carbonate corrosion prior to foraminiferal test fragmentation

is the combination of weight and light reflectance measurements of planktonic foraminiferal

tests of the polar species Neogloboquadrina pachyderma (sin.). The method was developed

by Helmke and Bauch (2002) and is restricted to regions and time intervals where carbonate

preservation is generally good. Carbonate corrosion leads to changes in the surface structure

of the calcite crystals and has a profound influence on the reflectivity of foraminiferal tests.

An inverse relationship between light reflectence and weight was found. Application to

Nordic Sea sediments revealed better preservation during glacial periods, which is consistent

with higher deep sea [CO32-] for this time scale. However, the method is yet far from being

used for quantitative estimates.

2.3 Estimating coccolithophorid paleoproductivity

Coccolithophorids are major contributors to the biogenic carbonate content in deep-sea

sediments (Archer et al., 2000; Milliman, 1993; Westbroek et al., 1993). Recently, there has

been increased interest in utilizing the elemental and isotopic chemistry of coccoliths. The

chemistry of coccolith carbonate may record different information than that of foraminiferal

carbonate because coccolithophorids, unlike foraminifera, are primary producers. Knowledge

about their paleoproductivity is of major importance for e.g. rain ratio estimates and

δ13Calkenone-based paleobarometer reconstructions (for review: Laws et al., 2001).

One limitation in the use of coccolith carbonate for geochemical studies has been their very

small size and therefore the inability to seperate monospecific coccolith assemblages for

analysis. New techniques now permit separation of fractions whose carbonate is highly

dominated (>70 % and often >90 %) by a single coccolith species (Stoll and Ziveri, in press).

As with foraminifera, calcite produced by different species of coccolithophorids has different

minor element partitioning and oxygen and carbon isotope fractionations. Specific

examination now opens a new field for paleoceanography.

Appendix 98

2.3.1 Coccolith Sr/Ca and stable carbon and oxygen isotopes to infer growth rate and cell size

The Sr/Ca ratio of coccoliths has been recently proposed as a potential indicator of

past coccolithophorid growth rates. The hypothesis is based on correlations between Sr/Ca in

polyspecific coccolith samples and primary productivity, alkenone-estimated growth rates,

and CaCO3 rain rates in deep sediment traps (e.g. Stoll and Schrag, 2000) across the

Equatorial Pacific upwelling region. Subsequently, a number of culture studies have

investigated controls over Sr/Ca ratios in coccoliths of several species. For identical

temperature and media composition, Sr partitioning is linearly related to rates of calcite

production/cell (Stoll et al., 2001). Higher calcification per cell at higher growth rates

observed in light-limited cultures of Gephyrocapsa oceanica, Calcidiscus leptoporus and

Emiliania huxleyi cultures (Paasche, in press; Stoll et al., in press) may suggest that active

uptake and calcification become increasingly important at higher growth rates. If coccolith

Sr/Ca is a reliable indicator of coccolithophorid productivity, it provides an index of past

productivity directly recorded by a primary producer. Furthermore, productivity estimates

from coccolith Sr/Ca do not rely on conventional determinations of sediment accumulation

rates which are often imprecise.

With regard to oxygen and carbon isotope fractionations, culture studies indicate

different nonequilibrium effects for different species of coccolithophorids (Dudley et al.,

1986; Ziveri et al., 2000; Ziveri et al., in prep.). These nonequilibrium effects appear to reflect

changing ecological and physiological responses of the organisms. In light- and nutrient-

replete cultures, the non-equilibrium effects in δ18O correlate highly with cell division rates

across a range of species. At similar calcification temperature and media composition, the

growth rates of the most common living species, E. huxleyi and G. oceanica, are strong and

δ18O is 3 ‰ offset with respect to equilibrium composition. In contrast, species with low

growth rates such as Umbilicosphaera sibogae var. foliosa have a δ18O fractionation effect of

~-2 ‰. Systematic relationships were also found between the carbon and oxygen isotopic

composition of the coccolith calcite for each species and the surface area/volume ratio of the

cells, which determines the diffusive flux of CO2 available to the cell (Ziveri et al., in prep.).

Clearly, more work is needed to test the validity of this proxy in constraining

coccolithophorid growth rates. Nevertheless, we are encouraged that qualitative or

quantitative determination of past variations in species-specific algal growth rates may be

possible.

Appendix 99

2.4 Model results 2.4.1 Lysocline reconstruction

In a modeling study, Jansen et al. (subm.) investigate the relationship between

lysocline and saturation horizon. The relative positions of these properties in a modern

sediment profile are described in Figure A3. The model results demonstrate that a

combination of changing CaCO3 and Corg production in combination with an increase in the

remineralization depth of organic carbon may have decoupled the lysocline and saturation

horizon during the last glacial maximum (LGM). Instead of the conservative estimate of 20

µatm, changes in the marine carbonate pump thus may have been responsible for a greater

portion of the observed glacial/interglacial atmospheric pCO2 shift on the order of 30-50

µatm.

Figure A3. a) Calcium carbonate content in a modern sediment profile. Lysocline and calcite

compensation depth (CCD) are set to the depth where the sediment calcite content drops below 95% and 10%, respectively. Near to the saturation horizon (SH), the rain rate of CaCO3 exceeds the dissolution rate, explaining the position above the lysocline. The transition zone between lysocline and CCD is indicated by the yellow band. b) Sedimentary calcite saturation profile. The solid blue line denotes the calcite saturation in bottom water, while the broken blue line refers to pore water saturation. Note that above -~3.6 km, the sediment is more corrosive than the bottom water due to organic carbon remineralization. Below that depth, calcite dissolution becomes progressively stronger, resulting in a pore water saturation state that is higher than in bottom waters.

Appendix 100

Geological records suggest that during the last glacial, the Atlantic lysocline was 0.3-1

km shallower (Crowley, 1983; Curry and Lohmann, 1986) than today, while it was about 0.8

km deeper in the Pacific Ocean (Farrell and Prell, 1989). Assuming that the lysocline has not

changed its position relative to the saturation horizon, these changes roughly correspond to a

decrease in atmospheric pCO2 by approximately 20 µatm (Broecker et al., 2001). However,

ice core observations suggest a glacial/interglacial shift in atmospheric pCO2 of 80 µatm

(Neftel et al., 1982; Petit et al., 1999). To bring the two records into line, additional reduction

in atmospheric pCO2 can be brought about by decoupling the lysocline from the saturation

horizon, due to respiration-driven carbonate dissolution in the upper 10 cm of sediments, as

observed e.g. by Hales and Emerson (1996). By changing the amount of organic carbon

arriving at the seafloor, the amount of carbonate dissolution above the saturation horizon can

change dramatically, shoaling the lysocline relative to the saturation horizon.

In contrast to the model used by Archer and Maier-Reimer (1994), Jansen et al.

(subm.) do not consider the dissolution of CaCO3 in the water column, as until now, the

underlying mechanism of this proposed feature has not been found (Jansen and Wolf-

Gladrow, 2001; Jansen et al., 2002; Milliman et al., 1999). Although this might result in an

overestimation of the decoupling, it does not affect glacial/interglacial changes in the whole

ocean carbonate inventory. Archer and Maier-Reimer (1994) tested scenarios for a glacial

ocean where Corg production was three times as high and CaCO3 production 60 % lower than

at present. They concluded that such a variation, operated by a shift from calcareous to

siliceous organisms during glacial times, might have driven atmospheric pCO2 to glacial

values. More recently, Sigman et al. (1998) argued that an increase in respiration-driven

calcite dissolution has no significant effect on the decoupling of lysocline and saturation

horizon as increased shallow water dissolution of carbonates would deepen the lysocline due

to mass balance considerations. However, it is questionable whether carbonate production and

dissolution are balanced at all (Milliman, 1993).

In contrast to Sigman et al. (1998), Jansen et al. (subm.) demonstrate that a decoupling

of lysocline and saturation horizon is possible. However, respiration-driven dissolution can

only significantly influence the atmospheric pCO2 when the export ratio depends on primary

production. Otherwise, the amount of organic carbon reaching deep-sea sediments would be

too small. With fixed export ratios, variability in CaCO3 production has a greater influence on

atmospheric pCO2 than variability in organic carbon production has.

Organic carbon productivity has been assumed to have increased during the LGM in

the range of up to +100 % relative to today (Berger et al., 1989; Kumar et al., 1995; Paytan et

Appendix 101

al., 1996), while CaCO3 productivity ranged between –60 % and +60 % relative to modern

times (Archer et al., 2000; Broecker and Henderson, 1998; Kumar et al., 1995). These

estimates yield Corg: CaCO3 rain ratios that are comparable to modern rain ratios in high

productivity areas. Thus, glacial pCO2 levels of ~ 230-250 µatm pCO2 are achieved within the

assumed rain ratio ranges. The result of Archer and Maier-Reimer (1994), who found that the

glacial to interglacial shift in atmospheric pCO2 is completely explainable by a decoupling

mechanism could not be reproduced by Jansen et al. (subm.). Rather, their model indicates

that ~40-60 % of the glacial pCO2 reduction may be attributed to changes in the marine

carbonate pump.

3. Recommendations

The proxies and model results discussed above demonstrate that the ideal proxy-

relationship does not exist. The record of a target parameter may be compromised in several

ways. However, the influence of many of the interfering parameters could be corrected for, if

the nature of the interferences were known and other proxies applied to estimate the

magnitude of the specific error. To uncover possible interferences, further laboratory culture

experiments on recorder organisms such as foraminifera, coccolithophorids and corals are

required. The magnitude of these effects needs to be quantified through additional field data

and process modeling. Of equal importance is a better understanding of the biomineralization

mechanisms of the major groups that provide proxies (i.e. foraminifera, corals,

coccolithophorids and diatoms). Only with such an understanding it will be possible to

explain the deviations of proxy relationships from thermodynamic predictions and to verify

whether such deviations are constant or require to be corrected for. Finally, the effects of

dissolution and other diagenetic changes need to be studied in laboratory experiments and at

depth.

3.1 Multi-proxy approach

The general advantage of preservational proxies (except examination with scanning

electron microscopy) is their non-destructive character. The descriptive analysis of weight,

light reflectance and general intactness leaves foraminiferal samples untouched and allows

further chemical investigation. Apart from providing additional information on bottom water

carbonate saturation, knowledge about the preservation state can furthermore help to elucidate

possible interferences with dissolution effects on other chemical proxies like Mg/Ca. Adding

chemical analyses to the interpretation of preservational proxies may reduce their own

Appendix 102

uncertainties. This could be for instance the application of U/Ca and SO42-/CO3

2- to estimate

surface water carbonate chemistry and its possible influence on weight and size of planktonic

foraminifera.

Probably the most detailed knowledge on interferences exists in the field of chemical

proxies in biogenic carbonates. Several culture studies on living foraminifera led to the

establishment of empirical proxy-relationships and the discovery of some of their limitations

(Bemis et al., 2000; Bijma et al., 1999; Erez and Luz, 1982; Erez et al. 2001; Lea et al. 1995,

1999; Lea and Spero 1992, 1994; Mashiotta et al. 1997; Rink et al., 1998; Russell, 2001;

Sanyal et al., 2001; Sanyal et al., 1996; Spero et al., 1997; Zeebe, 1999). Supplementary

information to many of these proxies could be obtained by application of multi-element ICP-

MS analysis on the same sample (Lea and Martin, 1996). However, for many parameters a

reliable proxy does not even exist.

A crucial requirement to minimize uncertainties in the background information is

therefore the search for new proxies for parameters such as salinity, productivity or deep sea

carbonate chemistry. A first approach into this direction is given by the application of time-

of-flight secondary ion mass spectrometry (TOF-SIMS, see section 3.3) to foraminiferal shell

analysis. Furthermore, new recorders and proxies need to be found to reconstruct the

conditions below the calcite compensation depth, where carbonates are already dissolved and

no longer available. All in all, there is a clear need to combine available chemical and

preservational proxies in order to improve the quality of paleoceanographic reconstructions.

3.2 Laboratory and field verification

As evident from previous sections, the establishment of empirical relationships in

laboratory and field experiments and calcification studies are important steps towards the

mechanistic understanding of a proxy and its paleoceanographic robustness. Especially for the

recently developed proxies it would be desirable to investigate these in more detail. Important

aspects would be the U/Ca-calibration of the warm water species G. sacculifer and G. ruber

and to obtain better information on the temperature dependency of the already existing

relationships. The boron isotopic composition of planktonic foraminifera has recently been

suggested to be biased by temperature and partial dissolution (Wara et al., 2001). To question

this finding, a set of planktonic foraminifera from a dissolution experiment is currently being

analyzed. However, temperature effects will better be investigated in culture experiments,

where other parameters can be kept constant. The effects of temperature and Mg on SO42-

/CO32- are currently being studied.

Appendix 103

Although they are the most important recorders of deep water (carbonate) chemistry,

benthic foraminifera are poorly represented in terms of experimentally calibrated proxy-

relationships. Culture experiments are therefore required to expand our knowledge on already

established and new proxies like U/Ca and SO42-/CO3

2-. It would be interesting to investigate

whether the carbonate ion effect on δ13C and δ18O is applicable to benthic foraminifera as

well.

Culture experiments are an important tool but reveal a rather restricted picture of a

proxy. Under natural conditions, planktonic foraminifera migrate vertically during their

ontogeny and finally sink towards the seafloor, thereby experiencing variable water column

conditions that modify their shell chemistry (Erez and Honjo, 1981; Lohmann, 1995). To

further open this black box, plankton tow samples, sediment trap material, and surface

sediment samples are needed to quantify the effects of this modification. Furthermore,

dissolution experiments under simulated natural conditions are suggested to elucidate if part

of the CaCO3 is dissolved preferentially, which proxies are affected by dissolution and to

what extent.

With regard to possible contaminating phases, the SO42-/CO3

2- proxy needs to be

investigated in the light of special cleaning efforts. For foraminiferal Ba/Ca analyses, barite

contamination was found to be removeable by cleaning with an alkaline DTPA

(diethylenetriamine-pentaacetic acid) treatment (Lea and Boyle, 1989). Cleaning experiments

may help to determine the robustness of this new proxy.

3.3 Recorders for deep water carbonate chemistry

As indicated in section 2.4.1, an increase of the marine carbonate pump under glacial

conditions may have led to increased respiration-driven carbonate dissolution in the sediment.

When arguing with sediment dissolution, the related diffusion of [CO32-] out of the sediment

into the overlying bottom water needs to be considered to possibly having biased the glacial

record, favouring locally more alkaline conditions than actually present in the deep water. At

least for the Holocene this limitation can be tested. For instance the difference in δ11B

between G. sacculifer and benthic foraminifera appears to be similar in magnitude to the

expected difference in pH between surface and deep water (Sanyal et al., 1995). However,

how can we be sure that benthic foraminifera, that are living right at the sediment-water

interface, always recorded deep water carbonate chemistry and not an anomalous local

environment? In fact, recent data by Anderson and Archer (2002) argue against the strong pH

increase suggested by Sanyal et al. (1995). To solve this problem, epibenthic deep water

Appendix 104

recorders are needed. For instance the chemical composition of the (less abundant) deep sea

corals, ostracod shells or fish otoliths has hardly been investigated but these skeletal parts

may possibly offer useful information with regard to deep and midwater (carbonate)

chemistry.

3.4 Development of new methods

One possibility to reduce the amount of carbonate required for boron isotopic analyses

may be the application of TOF-SIMS. This technique uses a focussed, energetic ion beam that

detaches particles and ions from a sample surface. Furthermore, high resolution images can be

obtained by scanning the surface of a sample with the beam of a liquid metal ion gun. The

combination of surface imaging with depth profile capabilities allows the visualization of the

three-dimensional distribution of elements in a foraminiferal shell. The overall advantage of

this method is the simultaneous detection of all elements and masses on very small samples

(single foraminiferal shells or less). The development of this method is still in progress. First

results were published by Crone et al. (2000; subm.) and Vering et al. (2001). In addition to a

substantial reduction of the amount of carbonate required for boron isotopic analyses, this

analytical technique could simplify the establishment of new proxies. However, each analysis

takes many hours to establish acceptable levels of precision, so that it is still far from

becoming a routine analytical tool.

3.5 Models 3.5.1 Thermodynamics of uranium uptake

In order to better understand the interaction between vital effects of a foraminifer and

sea water chemistry, D.A. Wolf-Gladrow and A.D. Russell intend to develop a diffusion-

reaction model comparable to Wolf-Gladrow et al. (1997) and Zeebe et al. (1999).

3.5.2 Lysocline reconstruction

The existing box model of Jansen et al. (subm.) is a first approach to describe the

processes that are important at certain sediment depth intervals. However, lysocline and

saturation horizon shifted in opposite directions in the Atlantic relative to the Indopacific

under glacial conditions (see section 2.4.1). In order to better estimate the global validity, the

model is intended to be coupled to a global circulation model.

Appendix 105

3.5.3 Boron isotopic fractionation in seawater

R.E. Zeebe is currently recalculating the stable boron isotope fractionation factor

between B(OH)4- and B(OH)3. This value has never been measured and is fundamental to the

boron isotope paleo-pH recorder.

Acknowledgements

The authors thank J. Bijma and G. Ganssen for organization and invitations to the

workshop. Financial support was provided by the European Science Foundation and the

Hanse Wissenschafts Kolleg.

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