*4.1. Brecciation and Carbonate Crystallization*

The 20 m-thick layer of tectono-sedimentary breccias drilled at Site 1277 (Figure 1; [37]) contains several calcite generations recognized by their crosscutting relationship and angular to rounded clasts of serpentinite (Figure 3). The calcite occurs in three contexts: it precipitates in veins crosscutting the clasts and formed during tectonic brecciation (type 1; Figure 4a,b); it constitutes the sedimentary component of the breccia cement (type 2; Figures 4b and 5b) and it crystallizes in the core of the serpentinite clasts (type 3; Figure 5a,b; [13]).

**Figure 3.** ODP Leg 210 Site 1277 4R-1 77–81 cm sample. Angular to rounded clasts of serpentinite are surrounded by a carbonate cement.

**Figure 4.** ODP Leg 210 Site 1277 4R-1 77–81 cm sample. (**a**) Calcite vein in serpentinite clast and (**b**) its cathodoluminescence equivalent (light orange corresponds to a high Mn concentration whereas dark red is associated with low Mn concentrations). (**c**) Microprobe map of MnO of the same area. (**d**) Microprobe profile of CaO, MgO and MnO versus the color intensity along the profile A, B shown by the red line in c. White crosses: data points of measured δ 18OSMOW and δ 13CVPDB.

**Figure 5.** ODP Leg 210 Site 1277 4R-1 77–81 cm sample. (**a**) Serpentine core mesh replaced by carbonate and (**b**) its cathodoluminescence equivalent (light orange corresponds to a high Mn concentration whereas dark red is associated with low Mn concentrations). White crosses: data points of measured δ 18OSMOW and δ 13CVPDB. (**c**,**d**) SEM images of calcite replacing serpentine.

In cathodoluminescence (CL) images, the calcite veins (type 1) display a color banding with an alternation of black and orange (Figure 4b). The banding is either parallel to fracturing, showing the opening direction of the vein (Figure 4b), or concentric, leading to scalenohedral crystals (Figure 5b). New grains appear to have nucleated in the vein at fracture edges, leading to single crystals with irregular shapes and low Mn-content (Figure 4d; dark in CL image). In addition, calcite grain growth occurs in parallel to the fracture opening direction (Figure 4b). It is common to observe in CL that the cores of the scalenohedron-shaped calcites are black (Mn-poor) whereas the last band is the brightest (Mn-rich; up to 1 wt. %, Figure 4d). The scalenohedral sparitic calcite grains found in veins vary in size from 10 to 100 μm. The banding in the veins display sharp contacts with the calcite in the cement due to brecciation (Figure 4c). The cement (type 2) is characterized by small, grain-sized calcite (<50 μm), with no CL-banding; CL colours are dark orange, with neither variation at sample or core scale (Figures 4b and 5b).

The serpentinite clasts are made of a network of mesh-like, regularly spaced fractures with two perpendicular orientations. Within the core of this regular mesh texture, 50 to 200 μm wide calcite grains are found (type 3) (Figure 5a,b). The mesh texture is preserved during replacement by carbonates (i.e., pseudomorphic replacement). In the core of the mesh, calcite grains grow concentrically from a nucleus, until reaching the rim of the mesh (Figure 5b). Depending on the nucleus location, the bands may not be continuous in the calcite grain. The CL colour of the bands is dark near the nucleus and again, bright orange in the outermost one (Figure 5b).

The banding observed in calcites of types 1 and 3 is similar (Figures 4b and 5b). At the contact between calcite and serpentine, the calcite grains are sharp and angular above the μm-scale (Figure 5c,d), and below the μm-scale the grains can display lobate shapes and are concave towards the serpentine (Figure 5d). The reaction front separates the well-crystallized calcite and fibrous serpentine and spans several micrometers. It is composed of <0.5 μm-thick calcite needles pervasively distributed in a disordered serpentine matrix (Figure 5d).

The μm-scale banding does not allow us to perform bulk isotopic analysis, as we would have obtained a mean value for all the bands. Therefore, we measured in-situ δ18O in the different bands with the SIMS. We acquired profiles through calcite grains in veins (type 1) and in the serpentine mesh core (type 3). Only CL bands larger than 20 μm could be analyzed (white crosses on Figures 4b and 5b; Table S2). Measured δ18OVSMOW values cluster in a relative restricted range for pseudomorphous grains (30.8 ± 0.4% to 32.6 ± 0.4%, *n* = 14), whereas calcite in veins are homogeneous (28.4 ± 0.4%, *n* = 49) (Figure 6; Table S2). In all the calcites measured, no variation of δ18OVSMOW can be observed from core to rim. <sup>δ</sup>13CVPDB measurements vary from 1.09 <sup>±</sup> 0.63% to 3.07 <sup>±</sup> 0.85% (Table S2) but do not show systematic variation with the Mn-banding. Our dataset sits within the range of measurements available for the Iberia margin (Figure 6).

**Figure 6.** Diagram of δ 18OSMOW versus δ 13CVPDB measured with the Secondary Ion Mass Spectrometer (SIMS) facility compared to compilation of literature data of ophicarbonates from the Newfoundland-Iberia (NF-I; squares) margin, Mid-Atlantic Ridge (MAR; circles) and Alpine ophiolites, affected by low metamorphism degree (black circles) and high metamorphism degree (black triangles). See references in Table S1.

#### *4.2. Thermodynamic Modeling of Carbonation*

Opposite flow directions in the recharge and the discharge models are responsible for the formation of different mineralogical assemblages (Figures 7 and 8) and aqueous fluid compositions (Figures 9 and 10). This is due to the variations in mineral solubility with temperature. If mineral solubility increases with temperature (silicates), fluid transfer from low to high temperature (recharge) will lead to dissolution, whereas fluid transfer from high to low temperature (discharge) will lead to precipitation. Therefore, mass transfer will differ for models considering recharge or discharge. An opposite trend is

expected for minerals with solubility decreasing with temperature (carbonates). Rock composition is observed to change along sharp boundaries (reaction fronts) for both recharge (Figure 7) and discharge models (Figure 8). In both models, the initial serpentine, exclusively composing the rock, is progressively replaced by talc as the fluid circulates through the rock. The talc is ultimately replaced by quartz (or clays at temperatures below ca. 40 ◦C) for the largest amount of fluid having circulated through the rock. The mineralogical assemblage can thus become quartz + carbonate and the modelled rock is a listvenite. Talc appears for lower (W/R)d when the precipitation of Mg-carbonates is allowed in the model, than when Mg-carbonates precipitation is excluded (Figures 7 and 8). These mineralogical evolutions require that the Mg and Si initially contained in serpentine are transported outside the rock as aqueous species in the fluid (Figures 9 and 10).

Carbonates are only formed at temperatures below 150 ◦C for the discharge model, and 100 ◦C for the recharge model (Figures 7 and 8e–h). In the discharge model, Ca concentration in the fluid slightly increases from 1.1 <sup>×</sup> <sup>10</sup>−<sup>4</sup> to 4.5 <sup>×</sup> <sup>10</sup>−<sup>4</sup> mol/kg as temperature decreases. The limitation of the carbonate stability field is thus mainly associated with pH decrease from 8.5 to 5 as temperature increases from 4 to 250 ◦C (Figure 10). In the recharge model, the limitation of the carbonate stability field is associated with anhydrite (CaSO4) formation at temperature above 100 ◦C as shown, for example in Figure 9a, with the simultaneous increase in Ca and decrease in HCO3 − concentrations at ~70 ◦C. Anhydrite formation only occurs in the first box of the model at high temperature in the discharge model (250 ◦C).

The distribution and composition of carbonates are different, even though the reactions of serpentine alteration were similar for discharge and recharge. The carbonates are first formed at the lowest temperature (4 ◦C) in the recharge model (Figure 7b,f). As fluid/rock interactions increase, the precipitation front progressively migrates towards higher temperatures. This front forming a serpentine + carbonate assemblage in equilibrium with a fluid of low Si concentration (Figure 9) is followed by a second front of talc + carbonate precipitation leading to increase in carbonate modal contents up to 60 mol. % (Figure 7h) and in Si concentration up to one order of magnitude (Figure 9h). Carbonates first replace serpentine at high temperatures (~150 ◦C) in the discharge model compared to the recharge model where this replacement starts at the lowest temperature (4 ◦C). The modal amounts of carbonate in the boxes are found to progressively decrease along the fluid flow path by one order of magnitude from 150 ◦C to 50 ◦C. Serpentine replacement by carbonates requires larger amount of fluids (or larger porosities for the same number of time steps in our model) to be observed in the case of recharge than in the case of discharge. This is due to the higher serpentine solubility where serpentine dissolution occurs in the discharge model at 150 ◦C than in the recharge at 20 ◦C. When Mg-carbonate precipitation is included in the model, carbonates are dolomite in the whole investigated temperature range in the recharge model and at temperatures below 20 ◦C in the discharge model. At temperatures above 20 ◦C, carbonates precipitate as magnesite in the discharge model, suggesting that Ca-rich carbonates are preferentially formed at low temperature. When Mg-carbonate precipitation is excluded, Ca-carbonates are only found to precipitate in the recharge model at temperatures below 110 ◦C (Figure 7).

**Figure 7.** Serpentine (green), carbonates (blue), talc (black) and anhydrite (yellow) modes as a function of temperature in the model for recharge (flow direction is indicated). The model on the left (**a**–**d**) does not include Mg-carbonates whereas the model on the right does (**e**–**h**). For each model, the modes are displayed at various dynamic water to rock ratios (W/R)d. Note that carbonates are first produced at low temperature in a reaction zone progressively extending towards higher temperatures. Carbonate formation is not predicted at temperature above 100 ◦C where anhydrite is the main Ca-bearing mineral to precipitate.

**Figure 8.** Serpentine (green), carbonates (blue) and talc (black) modes as a function of temperature in the model for discharge (flow direction is indicated). The model on the left (**a**–**d**) does not include Mg-carbonates whereas the model on the right does (**e**–**h**). For each model, the modes are displayed at various dynamic water to rock ratios (W/R)d. Note that Mg-carbonates are mainly produced at temperatures below 150 ◦C in the model allowing for Mg-carbonate precipitation whereas carbonates are not produced in the model in which Mg-carbonate precipitation is not allowed.

**Figure 9.** Si (black), Mg (green), Ca (light blue) and HCO3 − (dark blue) concentrations in the fluid and pH (dashed line) as a function of temperature in the model for recharge (flow direction is indicated). The model on the left (**a**–**d**) does not include Mg-carbonates whereas the model on the right does (**e**–**h**). For each model, concentrations and pH are displayed at various dynamic water to rock ratios (W/R)d.

**Figure 10.** Si (black), Mg (green), Ca (light blue) and HCO3 − (dark blue) concentrations in the fluid and pH (dashed line) as a function of temperature in the model for discharge (flow direction is indicated). The model on the left (**a**–**d**) does not include Mg-carbonates whereas the model on the right does (**e**–**h**). For each model, concentrations and pH are displayed at various dynamic water to rock ratios (W/R)d.

#### **5. Discussion**

#### *5.1. Serpentine Replacement by Carbonate During Seawater Influx in the NF Margin*

We identify three mechanisms of carbonation in the Newfoundland serpentinites: veining, cementation and pseudomorphic replacement. Tectonic deformation during mantle exhumation along passive rifted margins leads to cataclasis in the footwall of detachment faults [32] and to fluid circulation. Both mechanisms result in calcite vein formation (e.g., [5,60]). They follow the pre-existing fracture network of the serpentinized peridotite. Near the seafloor, sedimentation is active, leading to calcite cementation into serpentinite breccia. All along this tectono-sedimentary sequence, serpentine is partially replaced by carbonate in the core of the mesh texture. To determine when replacive carbonation occurs during the mantle exhumation, we measured the temperature record of calcite in the inferred oldest carbonate.

During mantle exhumation, the anisotropic thermal contraction of peridotite together with tectonic stresses generates a primary microfracturing responsible for the onset of serpentinization [61]. This fracture network is then re-used for reaction induced fracturing [56,62,63]. The serpentine growing along the microfractures is organized in lizardite pseudocolumns constituting the rim of the mesh texture [61,63,64]. For higher reaction degree, preserved olivine grains can be altered into isotropic serpentine or proto-serpentine [65,66]. The core of meshes can be composed of either preserved olivine or proto-serpentine, depending on the extent of the serpentinization reaction. Residual olivine grains are extremely rare in the Iberia and Newfoundland exhumed serpentine basement, where the degree of serpentinization is >90% complete [67]. Therefore, the different habits of serpentine crystals and their coherence is probably responsible for the preferential carbonate growth in the mesh core of the studied sample. The lobate shape of calcite grains suggests serpentine replacement. Our interpretation is reinforced by the fact that calcite needles at the reaction front start to grow from isotropic serpentine. In the veins, calcite grains grow with a similar CL banding, but the absence of lobate shape suggests that these grains are not replacive.

Hydrothermal systems are an important source of manganese as described in several present-day active hydrothermal vents, e.g., East Pacific Rise, [68] or in the acidic hydrothermal systems (pH < 5) from the MAR where Mn concentration varies from 59 to 2250 μM [18]. Ca2<sup>+</sup> from calcite can be substituted by Mn2<sup>+</sup> to form a solid solution of MnCO3 and CaCO3 [69]. CL of calcite is activated by Mn2+, whose emission colour is orange [70]. Temperature is not strongly influencing the Mn uptake in calcite grains [71]. Therefore, the Eh-dependent solubility of Mn into the fluid allows using CL colour variation as a proxy for oxygen fugacity variations in the fluid [72].

The CL-banding sequence is comparable between calcite grains growing in veins and in the serpentine mesh core, indicating a synchronous growth. The low content of Mn in calcite grains core reflects oxidizing conditions in both cases. In contrast, the last band in calcite is enriched in Mn testifying for more reduced fluids [72]. The fluid responsible for carbonation is oxidizing from the beginning of the reaction, even though fluids formed during serpentinization are reduced due to the coupled oxidation of the ferrous iron from olivine and pyroxenes to form magnetite [73]. This indicates that carbonation, from its onset, occurs after serpentinization near the seafloor, through interaction with oxidizing seawater.

The measured <sup>δ</sup>18O are homogeneous in calcite precipitating in the veins (δ18OVSMOW <sup>≈</sup> 31%) and vary by less than 2% in the calcite formed in the cores of the serpentine meshes. There is no systematic variation in the δ18O with respect to the Mn content variation or the location in the grain. The narrow ranges of the in-situ isotopic measurements at the tens of micrometer scale indicate limited variations of the composition of infiltrating fluids, both spatially, as suggested by the bulk data at MORs, and temporally, as indicated by the absence of zoning. Assuming seawater isotopic composition, the δ18O measurements give carbonation temperatures between Tmin −1 to 11 ◦C and Tmax 4 to 18 ◦C (Table S2). This low temperature of carbonate formation is consistent with experiments of inorganic carbonate precipitation at 25 ◦C [74–76]. It also suggests that serpentinization and carbonation can be two

temporally decoupled processes. Indeed, the highest serpentinization rates measured in experiments are found in the range of 250–300 ◦C [77,78].

The temporal decoupling between serpentinization and carbonation can also be shown by subtracting δ18Ocalcite from δ18Oserpentine [79]. For this calculation, we used a mean value calculated from a compilation of δ18Oserpentine from the Iberia margin (Leg 103, [34]; Leg 149, [3]) and from MAR [80]. The mean value for δ18Oserpentine used here is 10.65%. We also used a mean value for δ18Ocalcite from the NF margin i.e., 31%. If calcite was formed in equilibrium with serpentine, <sup>δ</sup>18Ocalcite–δ18Oserpentine <sup>≈</sup> 10%. Here, the difference exceeds 20% (Figure 6 and Table S2), therefore, calcite and serpentine are not at equilibrium. Fast advective fluid transport (e.g., in a fracture network) during fluid recharge or discharge would lead to limited interaction between the fluid and the serpentine, and to the preservation of the fluid isotopic signature. However, pervasive fluid flow in a low permeable rock favors the fluid–serpentine interactions and can thus modify the initial isotopic signature of the fluid and of the precipitating carbonate [81]. In addition to the fluid–serpentine disequilibrium, the extensive carbonation as well as the absence of δ18O zoning in calcite indicate that carbonation is fluid-buffered. As the fluid has a seawater signature, carbonation probably occurs near the surface during the onset of recharge into the serpentinized peridotites. The reproducibility of data in the MAR and Iberia-Newfoundland margins suggests that the recharge is a common process.

δ13CVPDB measured in NF margin are in agreement with previous data from the conjugated Iberia margin (i.e., between −3% and 3%) and are typical for marine hydrothermal systems (e.g., [1]). This suggests that C is dominantly inorganic derived from seawater as suggested for the Iberian margin [1]. Thermodynamic modeling indicates that such a precipitation of seawater-derived inorganic carbon is possible and requires the circulation of more than 10<sup>3</sup> kg of seawater per kg of rock (Figures 7 and 8). The need for high dynamic water to rock ratios is associated with the low concentration of inorganic carbon in seawater (~2 mM). There are no systematic correlations between Mn-banding and δ13C variation.

#### *5.2. Carbonation in Passive Margins and Slow-Spreading MORs*

We compiled calcite δ18O and δ13C bulk measurements from the Iberia and Newfoundland margins and from the Central MAR ophicarbonates (Figure 6). Based on their δ18O record and calculated temperatures of formation, we infer that the ophicarbonates can be separated into two distinct groups, (1) calcite formed during seawater influx in an evolving carbonation front; or (2) during hydrothermal fluid discharge, in the case that during recharge, fluid-peridotite interaction was small.

As calcites from the Iberia margin and MAR, formed by veining, cementation and pseudomorphic replacement, record high δ18O, we infer this calcite to be formed during seawater influx. This has already been proposed in calcite veins from Iberia with O isotopic compositions varying from 27.6% to 31.2% [1] and in sedimentary carbonates (e.g., botryoidal calcite) at the Iberia margin (δ18O from 30% to 32%; [2,35]) and in the MAR (δ18O from 33% to 35%; [23]).

Calcite formed during hydrothermal fluid discharge is created by veining or hydrothermal deposits (e.g., [21,23,24]). In the MAR, carbonates found in veins and deposits are calcite and aragonite [21]. Mg content in calcite can be as high as 28 wt. % [24]. These high temperature calcite veins are found at the vicinity of hydrothermal vents, and record δ18O between 26% and 8% (Mid-Atlantic Ridge, Kane Fracture Zone [26]; ODP Leg 209, [27]; Lost City, [21,23]). However, those samples represent a minority of the collected samples (see references in Table S1). Such low δ18Ocalcite indicates carbonation temperatures higher than 100 ◦C [23]. No pseudomorphic replacement is observed in environments in close relationship with an active hydrothermal vent. Based on O and C isotopic signature of replacive calcites, we interpret carbonate replacement in the samples from Newfoundland, the Iberian margin and the MAR as related to near-surface alteration during seawater influx.

We demonstrate that carbonation on passive margins is a low temperature, near surface process. This is true, even if magmatism is present as demonstrated at Site 1277, where syn- and post-rifting magmatism was identified (e.g., [39]). Indeed, calcite recorded temperatures below 20 ◦C. In the case

of the Iberia-Newfoundland margins, magmatism has also no effect on the isotopic record, and hence does not influence carbonate precipitation.

#### *5.3. Insights from Numerical Modeling*

Carbonates are formed at two distinct temperature ranges at MORs, either below 10 ◦C or locally at the Lost City above 100 ◦C in veins feeding hydrothermal vents. Interestingly, this difference in carbonation temperature is also the main difference between the two simulations performed here with recharge and discharge models. Our simulations reproduced carbonate formation during recharge at near-surface conditions. They also predict carbonate formation at approximately 150 ◦C during discharge. Moreover, carbonate production required a higher fluid amount ((W/R)d) for recharge than for discharge models, suggesting more open system conditions. Talc formation is decoupled from carbonate precipitation in the recharge models only, leading to rocks exclusively composed of serpentine and carbonates as observed in the natural samples. In all the simulations, Ca-rich carbonates were only formed at low temperature, whereas carbonation at temperature above 150 ◦C only produced Mg-rich carbonates. Based on these results, we interpret carbonate replacement in the samples from Newfoundland, the Iberian margin and the MAR as related to near-surface alteration during recharge. Carbonation at high temperature (>100 ◦C; [23]) is also expected to occur at MORs during discharge and, in particular, near hydrothermal vents such as the Lost City. The simulations predict the observation of calcite in the natural samples studied here, but they do also predict Ca-rich carbonates formation at temperatures above 100 ◦C, as is observed at the Lost City Hydrothermal Field. Due to missing compatible thermodynamic data, we do not include phases such as layered double hydroxides (coalingite-pyroaugite, LDHs), and hydrous Mg-carbonates (hydromagnesite and nesquehonite) (e.g., [82]).
