GSA Data Repository 2016188

The control of silicate weathering by interface coupled dissolution-

precipitation processes at the mineral-solution interface

Encarnación Ruiz-Agudo (1)*, Helen E. King(2), Luis D. Patiño-López(3), Christine V. Putnis(4,5),

Thorsten Geisler (6), Carlos Rodriguez-Navarro (1) and Andrew Putnis (4,7)

(1) Departamento de Mineralogía y Petrología, Universidad de Granada, 18071 Granada,

Spain

(2) Department of Earth Sciences, Utrecht University, 3584 CD Utrecht, The Netherlands

(3) Centro de Investigación Científica de Yucatán, 97302 Mérida, México

(4) Institut für Mineralogie, Universität Münster, 48149 Münster, Germany

(5) Nanochemistry Research Institute, Department of Chemistry, Curtin University, Perth 6102,

Australia

(6) Steinmann Institut für Geologie, Mineralogie und Paläontologie, University of Bonn. 53115

Bonn, Germany

(7) The Institute for Geoscience Research (TIGeR), Curtin University, Perth 6102, Australia

(*) corresponding author: encaruiz@ugr.es

SUPPLEMENTARY INFORMATION

DR1. MATERIALS AND METHODS.

- Materials. High purity wollastonite crystals from Barberton District, Mpumalanga

Province, (South Africa) were cleaved with a knife blade to obtain mm-size fragments and used

in the dissolution experiments. Solutions were prepared immediately before the experiments

using double-deionized water (resistivity >18 m cm-1) and adjusting the solution pH down to

pH 1.5 using HCl. The absence of calcium and silicon in the input solutions ensured initial far-

from-equilibrium conditions with respect to the original wollastonite.

- Ca and pH measurements using microelectrodes. Batch dissolution experiments were

performed in open Teflon® reactors of 30 mL inner volume. Wollastonite crystals of ca. 8 mg

were glued to the bottom of the reactor using carbon glue before contact with 30 mL of stagnant

or flowing (60 mL h-1) acidic aqueous solution (HCl pH 1.5). Microelectrodes were used for

on-line measurement of pH (micro combination electrode mod. PHR-146B, Lazar Laboratories)

and calcium concentration, [Ca] (micro-Ion Selective Electrode mod. LIS-146CACM with

microdouble junction reference electrode mod. DDM-146, Lazar laboratories) during

dissolution. The detection limit of the Ca probe is 1 ppm, and the instrumental error is ± 2%.

Electrodes were placed at different distances (0 to 3.8 mm) from the mineral surface using

home-built equipment. Free-Ca and pH data were collected at 0, 254 and 653 m from the

mineral surface and subsequently at 653 m intervals. A mechanical step-displacement

micrometer (Model 9664-0107, Limit, Sweden) was used to set the height of the microprobes.

The response time of the pH electrode is 1 second, and the time in each position (total number

of positions = 6) was 10 seconds. Therefore, at the time in which the reading was recorded, the

probe had had enough time to settle in the corresponding position, reporting a constant value.

Measurements in each point were carried out every 60 seconds. Once the data were collected,

going up and down in distance from the mineral surface with the time, the plot was constructed

by interpolation of the information between measured points using Microcal Origin 6.0

software. Microelectrodes were calibrated using standard pH buffered solutions and CaCl2

solutions of known concentration (0.1 to 100 mM). Experiments were repeated up to three times

to test the reproducibility of the results.

The time evolution of the maximum saturation index, SI (SI = log(IAP/Ksp), where IAP

and Ksp are the ion activity product and solubility product of a relevant phase, respectively) with

respect to wollastonite and amorphous silica was calculated using PHREEQC (Parkhust and

Appello, 1996) for the different values of pH and [Ca] (assuming stoichiometric release of Ca

and Si) with respect to wollastonite and amorphous silica. All calculations were performed

considering STP conditions. For the calculation of SI, we used Ksp values of amorphous silica

and wollastonite in the minteq.dat database (-3.0 and 13.0 at 25 ºC, respectively), for the

following dissociation reactions:

SiO2 + 2H2O = H4SiO4 (1)

CaSiO3 + H2O + 2H+ = H4SiO4 + Ca+2 (2)

Because of the lack of direct Si concentration measurements (micro-Si probes are not

available on the market), the observation of regular etch pits forming and spreading (Ruiz-

Agudo et al. 2012) is critical to demonstrate that Ca and Si are released stoichiometrically

(otherwise etch pits could not expand regularly). However, big silicate units cannot diffuse in

the fluid as fast as Ca2+. Note that the concentration gradients calculated here represent limit

(lowest) values; differences in diffusion rates between silica species and calcium would only

result in higher Si concentrations and SI with respect to amorphous silica (and thus silica would

be more likely to precipitate). Therefore, even without quantifying this effect, the conclusions

are fully valid and our calculations can be considered as conservative.

No other phases are considered since experimental studies have shown that surface

altered layers formed upon incongruent dissolution of silicate minerals correspond to

amorphous silica (e.g., Casey et al. 1993; Ruiz-Agudo et al. 2012).

Figure DR1. Time evolution of solution chemistry in flow-through tests. Contour maps

showing the time evolution of the solution (a) free calcium concentration (mM), (b) pH and

saturation index, SI, with respect to (c) wollastonite and (d) amorphous silica during

wollastonite (flow rate 60 mL h-1) dissolution experiments in an acidic aqueous solution (pH

1.5, HCl).

Figure DR2. Time evolution of solution chemistry in unstirred tests. (a), (b) and (c) show

amorphous silica saturation index profiles at 200, 500 and 1000 seconds, respectively.

After completion of dissolution tests, wollastonite crystals were recovered, rinsed with

distilled water and dried at room T prior to mounting on Al stubs. Samples were subsequently

C-coated and observed and analyzed under a field emission scanning electron microscope

(FESEM, Zeiss Supra 40VP) equipped with EDS microanalysis (Figure DR2).

 

Figure DR3. Textural and compositional features of wollastonite subjected to flow-through

dissolution test. a) FESEM photomicrograph of wollastonite crystal displaying the (100) face

and normal to it, a rough, stepped (hkl) face where SAL is developed; b) detail of SAL from the

squared area in (a) showing lamellar dissolution features forming SAL and pyramidal-shaped

areas without SAL (lower part of the image). c) EDS compositional AB profile in (b) showing

Si-enriched zones corresponding to SAL. Note that some areas display a composition consistent

with unaltered wollastonite, demostrating that the SAL does not cover the full surface of

wollastonite.

- 18O tracer experiments. Batch isotopic tracer experiments of amorphous silica

formation during wollastonite reaction in acidic solutions were run in polytetrafluoroethylene-

lined steel autoclave batch reactors at 90 °C for three days in 1 mL of 0.1 M HCl solutions. Two

experiments, one with water enriched in 18O and one with MilliQ water, i.e. natural 18O

concentrations hence 16O-rich, were run in parallel to examine the 18O uptake into the

amorphous silica. Both solutions were diluted from a 1 M HCl solution prepared using 37%

HCl and MilliQ double deionized water (resistivity >18 m cm-1). To produce the 18O enriched

solution 97 at % 18O-enriched water (Campy Scientific) was used to dilute the 1 M HCl solution

to 0.1 M and a similar enrichment (67 at %) to that used in O isotopic tracer experiments with

the isosilicate olivine (King et al. 2011). The solutions were then added immediately to a pre-

weighed wollastonite fragment (~0.2 g) in the polytetrafluoroethylene liner and sealed. The

autoclaves were weighed whilst sealed before the experiment, after 24 hours and upon

completion of the experiment to ensure that no solution had been lost. The oven was preheated

to 90 °C before the autoclave was placed inside. Upon termination of the experiment each

autoclave was quenched using pressurized air. After reaching room temperature the autoclaves

were opened and the solution removed. The remnant wollastonite grains were then placed on

filter paper to remove any remaining solution before being left to dry in air for 24 hours. When

dry the samples were embedded in epoxy and cross-sectioned for scanning electron microscopy

(SEM) and Raman spectroscopic analysis.

Measurements of isotope incorporation into the amorphous silica were conducted using

a Horiba Scientific LabRam HR800 confocal micro-Raman spectrometer at Bonn University.

Raman scattering was excited using a solid state Nd:YAG laser (532.09 nm) with an intensity of

about 150 mW at the surface. A 100x objective lens with numerical aperture of 0.9 provided a

laser spot size of less than 1 µm at the surface. The confocal hole was opened to 300 µm to limit

the depth penetration of the laser and therefore the amount of wollastonite sampled from the

underlying material. After passing through a 50 µm entrance slit and being dispersed by a

grating of 1800 grooves/mm the scattered Raman light was collected in 180° geometry by an

electron-multiplier charge-coupled device. The Raman microspectrometer was calibrated using

the 520.7 cm-1 band of a silicon wafer standard.

The rim material was identified as amorphous silica in both experiments based on the

development of the bands relating to siloxane ring structures below 485 cm-1. At least six

locations within the amorphous silica rim were examined for each of the samples to examine the

consistency of isotope incorporation. Similar to reports by Casey et al. (1993), wollastonite

peaks were also present in the majority of the Raman spectra. A few areas were observed with

minimal contamination from the wollastonite; however, due to the porous nature of the

amorphous silica these spectra also had contributions from the epoxy used to embed the samples

(e.g., Figure DR4a), which obstructed some of the amorphous silica bands. The Rayleigh

scattering contribution to the background in the low wavenumbers region was removed using

the procedure described by Long (1976) and recently used by Behrens et al. (2006) for the study

of amorphous silica. Final background removal for the entire spectrum was subsequently

completed using a spline cubic baseline. Further data reduction, such as removal of the epoxy

related spectrum, and peak fitting was conducted using the LabSpec5 program also from Horiba

Scientific.

All locations within the amorphous silica produced during the 18O-enriched experiments

showed a red shift in the D1 peak of 10 cm-1 in comparison to the 16O-rich experiment, which

was confirmed by the same shift of the SiO4-SiO4 peak (Figure DR4b). This red shift is

consistent with 18O incorporation into the amorphous silica structure. Si18O2, i.e. amorphous

silica with only 18O atoms, produces a frequency shift of -30 cm-1 for the D1 band (Galeener and

Mikkelsen, 1981). The principle behind analyzing the 18O content in condensed matter by

vibrational spectroscopy is that the energies or frequencies associated with vibrational motions

are dependent on the masses of the vibrating atoms (Herzberg 1944). In the simple harmonic

approximation, the vibrational frequency shift is proportional to the square root of the mass ratio

between the atoms involved in the vibration (Herzberg 1944). A complete substitution of 16O by

18O in a silica network was found to give a mass-related frequency shift of −30 cm-1, i.e., −0.3

cm-1/at.% 18O (Galeener and Mikkelsen 1981). The measured 10 cm-1 shift of the D1 band to

lower wavenumbers thus corresponds to ~30 at.% 18O within the amorphous silica structure.By

assuming a simple linear relationship between the 18O content and frequency shift we calculate

that the observed shift is caused by about 34 at% 18O within the amorphous silica structure. A

shift of up to 17 at% can be related to the hydroxylation of the silica chain during dissolution,

therefore the calculated enrichment indicates that silica was also released into solution where it

was free to exchange O-atoms with the surrounding water molecules. However, the calculated

enrichment is lower than that observed for the isosilicate olivine (King et al. 2011). One

possible explanation for this would be the release of silica as dimers or trimers during the silica

chain breakdown where bridging O atoms are not as freely exchangeable as the hydroxyl groups

present on monomer silicic acid.

Figure DR4. Raman spectra obtained from wollastonite reacted in 0.1 M HCl for 3 days

at 90 °C. (a) Sample spectra clearly shows presence of bands related to siloxane ring structures

in amorphous silica below 500 cm-1 that cannot be attributed to wollastonite or epoxy

contamination. (b) Raman spectra from amorphous silica produced during an experiment 18O-

enriched and 16O-rich solutions where the contributions from epoxy and wollastonite bands have

been removed.

- Mach-Zehnder phase-shift interferometry. Similar batch and flow-through

experiments were performed in reaction cells (2 mL volume) using a specifically built Mach-

Zehnder interferometer (Figure DR3). The instrument follows an unbalanced interferometer

scheme, providing phase shifting through laser diode (635 nm) wavelength variations (Ishii,

2004), and also a low spatial coherence illumination with an RGG diffusor for imaging

improvement (Dubois et Al, 2004). In Mach-Zehnder interferometry, a laser beam is divided

into the reference and the sample (test) beams by means of a beamsplitter. The sample beam

propagates through the experimental volume, and is locally delayed by concentration variations,

before being recombined with the reference beam. Interference fringes at the imaging plane

appear as a result of the recombination, and collected with a CCD camera. The so obtained

interferogram carry the information of the optical path variations.

Phase cannot be directly extracted from a single interferogram where pixel intensity is

proportional to the cosine of phase difference. Phase shifting interferometry (PSI) allows

extracting the real phase for each pixel. It consists in acquiring a minimum of 3 phase-shifted

images of the same fluid status and in calculating the phase of each pixel with a phase shifting

algorithm. The method consists in introducing in the reference beam a small (controlled) phase

delay between the recorded images.

In this work, a set of five phase-shifted images has been acquired and phase is retrieved

by means of Hariharan algorithm (Hariharan et Al, 1987). The phase map is obtained by means

of an expression involving the arctangent function, which returns values that are known

between the limits π and ‐π. Hence 2π and discontinuities with values near to 2π appear in the

phase distribution. Unwrapping is the procedure by which these discontinuities are resolved; the

result is converted into the desired continuous phase function. Phase was unwrapped according

to the algorithm developed by Herráez and co-workers (Herráez et al. 2002). Finally, a set of

several tenths of consecutive phase maps acquired just before the introduction of the crystal into

the cell was acquired and averaged, and the resulting phase map was employed as reference and

subtracted to each frame obtained during the dissolution process. Average Refractive Index

variations, Δn along the observation axis z are retrieved from the phase variations according to:

Δ , , , Δ , , (1)

Where is the laser wavelength and Z is the sample cell thickness along the observation

axis z.

Note that these experiments, although do not provide direct quantitative compositional

information, they allow a higher spatial and temporal resolution than the quantitative

measurements performed using Ca and pH microelectrodes. In particular, in the case of flow-

through experiments, although Ca, pH measurements and calculated saturation indexes indicate

that the solution remains slightly undersaturated with respect to amorphous silica, PSI

observations indicating periodic variations of solution composition and FESEM-EDS analysis

showing spatially and locally constrained formation of discontinuous SALs suggest that

supersaturation with respect to amorphous silica may have been achieved temporarily and in

restricted areas of the wollastonite surface during flow-though dissolution tests. Such variations

are not detected by compositional measurements performed using Ca and pH microelectrodes,

as they average compositional data in the influence area of the microelectrode and are collected

at ca. 240 s intervals.

Figure DR5. Mach-Zehnder interferometry set-up used in this study.

DR2. EFFECT OF SURFACE ALTERED LAYERS ON MODELING OF GLOBAL

TEMPERATURE AND ATMOSPHERIC CO2. Unlike climate models and simulations based

on a “top-down” approach, in which the implemented weathering equations are empirical, not

based on any fundamental mechanisms and activation energies are commonly taken from field

studies (e.g. GEOCARB, Berner and Kothavala, 2001), in calculations based on a “bottom-up”

approach, weathering is approximated by means of quantitative expressions, which mostly

depend on mechanisms and rates determined in the laboratory (Brady 1991). As in Brady

(1991), we will use Volk’s carbon cycle model (Volk, 1987) as a tool to quantify the influence

of surface altered layers (SALs) on global carbon cycling. This relatively simple model has been

shown to give essentially the same results as other computationally more complex model, such

as BLAG (Volk, 1987 in Brady 1991). In this model, the relationships

1 0.0381 1

2.88,

can be used to calculate atmospheric CO2 concentrations and paleotemperatures as a function of

Ea and the independently measured geophysical forcing parameters fsr and fa (Volk, 1987).

These two latter parameters provide the link between tectonics and climate. fsr/fa is taken from

Figure 3 of Brady (1991) (Figure DR6). T0 is set to 288.15 K, the preindustrial global mean

surface temperature. Ea is the activation energy for silicate weathering, and represents the

strength of the feedback between weathering and global temperature. Ca and Mg silicate

weathering is the primary sink for atmospheric CO2 and thus it seems logical to use

experimentally determined Ea in Ca and Mg silicate dissolution experiments to parameterize the

atmospheric CO2 concentrations and temperature dependency of weathering (Brady, 1991).

Equations 1 and 2 are solved for three different cases: (a) Ea=18.9 kcal mol-1, determined in

congruent wollastonite dissolution experiments (Rimstidt and Dove, 1986); (b) Ea= 13.1 kcal

mol-1 (Bailey 1977 as stated in Brantley and Chen 2004) and (c) Ea=11.2 kcal mol-1 (Ptácek et

al. 2011), both determined in “incongruent” dissolution experiments, with formation of surface

altered layers.

Figure DR6. Ratio of and seafloor spreading rates fsr to continental surface area fa over the past

90 m.y. Values of fa from Barron et al. [1980] and values of fsr from BLAG are taken from

Figure 3 of Brady (1991).

When low Ea (down to 10 kcal/mol) determined for incongruent dissolution of

wollastonite (with formation of surface-altered layers (Bailey, 1976; Brantley and Chen, 1995;

Ptáček et al. 2011) are input into existing models that calculate paleo-T and atmospheric CO2

levels over the past 90 Myr based on a mechanistic approach of mineral reactivity, the result is

higher CO2 levels and temperatures up to ca. 2 °C warmer than those estimated by assuming

higher Ea (20 kcal/mol) determined in congruent dissolution experiments (Rimstidt and Dove

1986) (Figure DR7a). Note that the low and high Ea values for wollastonite used in our

modeling bracket the values determined in the laboratory for common primary silicates such as

feldspar, olivine, amphiboles and other pyroxenes such as augite and diopside (Brantley 2008;

Brady 1991). They also fall within the range of Ea values (5.7-20.5 kcal/mol) determined in the

field for the dissolution of different silicate minerals and glasses (Velbel 1993), the lower values

suggesting a diffusion control. Despite its simplicity, this modeling yields a good match with

pCO2 levels determined using independent proxies (Beerling and Royer, 2011) (Fig. DR7b).

Interestingly, the best match is achieved when considering the low Ea value, that is, when

considering the case of wollastonite dissolution under transport-controlled kinetics and resulting

in the development of a SAL. These results illustrate the impact that different weathering

mechanisms and rate-controlling steps in laboratory experiments may have on estimates of

paleo-T and pCO2 levels in the past as well as in the prediction of future climate evolution using

models in which the kinetic rate laws for the weathering dependence on temperature are based

on a mechanistic approach of mineral reactivity. For instance, following a similar reasoning, the

long intervals between snowball glaciations (750-550 My), several orders of magnitude longer

than expected from typical rates of CO2 drawdown by silicate weathering (Mills et al. 2011),

could be explained by considering the effect of transport-limited silicate weathering on

atmospheric CO2 levels. Initial silicate weathering in a CO2-rich, acidic aqueous environment

would result in the quick coating of SALs on reactive rock-forming minerals. This would in turn

lead to a stage of transport-limited silicate weathering in which CO2 would be removed from the

atmosphere at a slower rate than in the case of reaction-limited weathering, reducing the rate of

atmospheric CO2 drawdown to threshold values needed for the beginning of a new glaciation.

Figure DR7. Effect of surface altered layers on modeled global temperature and

atmospheric CO2. (a) Maximum partial pressure of CO2 in the atmosphere (in parts per million

per volume, ppmv) and mean global surface temperature (inset) over the past 90 Myr, calculated

using Volk’s carbon cycle model50 and different activation energies, Ea, determined in

experimental studies: 18.9 kcal mol-1, determined in congruent wollastonite dissolution

experiments20; 13.1 kcal mol-1 25,40 and 11.2 kcal mol-1 41, both determined in “incongruent”

dissolution experiments, with formation of SAL. (b) Comparison of paleo-pCO2 values

determined by different proxies (from ref. 43) and estimated values in (a) over the past 65 Myr.

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