Lower crustal hydrothermal circulation at slow-spreading ridges: evidence from chlorine in Arctic and South Atlantic basalt glasses and melt inclusions

  • Froukje M. van der Zwan
  • Colin W. Devey
  • Thor H. Hansteen
  • Renat R. Almeev
  • Nico Augustin
  • Matthias Frische
  • Karsten M. Haase
  • Ali Basaham
  • Jonathan E. Snow
Original Paper


Hydrothermal circulation at slow-spreading ridges is important for cooling the newly formed lithosphere, but the depth to which it occurs is uncertain. Magmas which stagnate and partially crystallize during their rise from the mantle provide a means to constrain the depth of circulation because assimilation of hydrothermal fluids or hydrothermally altered country rock will raise their chlorine (Cl) contents. Here we present Cl concentrations in combination with chemical thermobarometry data on glassy basaltic rocks and melt inclusions from the Southern Mid-Atlantic Ridge (SMAR; ~ 3 cm year−1 full spreading rate) and the Gakkel Ridge (max. 1.5 cm year−1 full spreading rate) in order to define the depth and extent of chlorine contamination. Basaltic glasses show Cl-contents ranging from ca. 50–430 ppm and ca. 40–700 ppm for the SMAR and Gakkel Ridge, respectively, whereas SMAR melt inclusions contain between 20 and 460 ppm Cl. Compared to elements of similar mantle incompatibility (e.g. K, Nb), Cl-excess (Cl/Nb or Cl/K higher than normal mantle values) of up to 250 ppm in glasses and melt inclusions are found in 75% of the samples from both ridges. Cl-excess is interpreted to indicate assimilation of hydrothermal brines (as opposed to bulk altered rock or seawater) based on the large range of Cl/K ratios in samples showing a limited spread in H2O contents. Resorption and disequilibrium textures of olivine, plagioclase and clinopyroxene phenocrysts and an abundance of xenocrysts and gabbroic fragments in the SMAR lavas suggest multiple generations of crystallization and assimilation of hydrothermally altered rocks that contain these brines. Calculated pressures of last equilibration based on the major element compositions of melts cannot provide reliable estimates of the depths at which this crystallization/assimilation occurred as the assimilation negates the assumption of crystallization under equilibrium conditions implicit in such calculations. Clinopyroxene–melt thermobarometry on rare clinopyroxene phenocrysts present in the SMAR magmas yield lower crustal crystallization/assimilation depths (10–13 km in the segment containing clinopyroxene). The Cl-excesses in SMAR melt inclusions indicate that assimilation occurred before crystallization, while also homogeneous Cl in melts from Gakkel Ridge indicate Cl addition during magma chamber processes. Combined, these observations imply that hydrothermal circulation reaches the lower crust at slow-spreading ridges, and thereby promotes cooling of the lower crust. The generally lower Cl-excess at slow-spreading ridges (compared to fast-spreading ridges) is probably related to them having few if any permanent magma chambers. Magmas therefore do not fractionate as extensively in the crust, providing less heat for assimilation (on average, slow-spreading ridge magmas have higher Mg#), and hydrothermal systems are ephemeral, leading to lower total degrees of crustal alteration and more variation in the amount of Cl contamination. Hydrothermal plumes and vent fields have samples in close vicinity that display Cl-excess, mostly of > 25 ppm, which thus can aid as a guide for the exploration of (active or extinct) hydrothermal vent fields on the axis.


Hydrothermal circulation (Ultra)slow-spreading ridges Crystallization depths Crustal assimilation MORB Chlorine 


When new oceanic lithosphere is formed at a spreading centre, heat is transferred upwards from the mantle either by advection (magmatism) or tectonism. Although there is fairly convincing evidence that hydrothermal circulation of seawater cools the upper parts of this system (e.g. Stein and Stein 1994), up to present no consensus on the maximum depth of hydrothermal circulation in newly formed oceanic crust (e.g. Mottl 2003) or what role hydrothermal circulation plays in the deep cooling of ridges has been reached. Particularly at (ultra)slow-spreading (< 1.2 cm year−1 full spreading rate) and slow-spreading ridges (< 1.2 to < 5.5 cm year−1) cooling of the lower lithosphere by hydrothermal circulation may be relevant, as seafloor is partially produced by tectonic processes (e.g. Dick et al. 2003; Cannat et al. 2006) and the lack of permanent magma chambers and low crustal temperatures (e.g. Harper 1985; Detrick et al. 1990) may allow faulting to be particularly extensive and to penetrate deeper, down to the lower crust, potentially providing pathways for hydrothermal fluids (Harper 1985; Cannat et al. 1991; Mével and Cannat 1991).

Direct evidence for deep circulation comes from studies of oceanic drill sections (from crust formed at fast-spreading ridges) and from ophiolites (e.g. Gregory and Taylor 1981; Stakes and Vanko 1986; Nehlig and Juteau 1988; Lecuyer and Reynard 1996; Kawahata et al. 2001; Coogan 2003; Nicolas et al. 2003), both of which display high-temperature alteration of the lower crust. Deep hydrothermal circulation at fast-spreading ridges is further indicated by thermodynamic modelling (Cherkaoui et al. 2003; Maclennan et al. 2005; Hasenclever et al. 2014) and the seismic structure of the lower crust (Dunn et al. 2000). Evidence from slow-spreading ridges include P-wave tomography underneath the TAG hydrothermal field, which indicates hydrothermal circulation, related to a detachment fault, extending at least 2 km, but potentially 3.5 km deep (Zhao et al. 2012). Other direct information on the depth of penetration of hydrothermal circulation at slow-spreading ridges is only available from drill-cores that have a limited depth extent (e.g. < 1600 m in Alt and Bach 2006) or from tectonically exposed parts on the lower crust (Gillis et al. 1993), which have thus been exposed to ambient seawater for extended periods of time.

Deep hydrothermal circulation could potentially be traced by detecting it effects on the chlorine budget of erupted magmas if those magmas assimilated the hydrothermal fluids themselves or hydrothermally altered country rock at depth (Michael and Schilling 1989; Michael and Cornell 1998; Kent et al. 1999; Coogan et al. 2003; Gillis et al. 2003; Sun et al. 2007; France et al. 2009; Kendrick et al. 2013; van der Zwan et al. 2015). Assimilation of surrounding rocks by rising MORB magmas has been shown to take place in several gabbroic complexes (e.g. Bédard et al. 2000 and references therein; Godard et al. 2009; Drouin et al. 2009; Sanfilippo et al. 2014; Fischer et al. 2016), but the depth and/or hydrothermal alteration state of the assimilated rocks is uncertain so far.

Although work on fast-spreading ridges indicates addition of Cl relative to trace elements of similar incompatibility (e.g. K or Nb) in a reaction zone between the magma and the overlying hydrothermal circulation cell at the roof of a magma chamber (Coogan et al. 2003; Gillis et al. 2003; le Roux et al. 2006; France et al. 2009; Kendrick et al. 2013), previous studies on slow-spreading ridges (e.g. Michael and Cornell 1998) had not found evidence of Cl addition to their magmas. Michael and Cornell (1998) argued that because fast-spreading ridges generally have shallow (100–300 MPa ≈ 3–10 km) axial magma chambers whereas magmas from a typical slow-spreading ridge (e.g. not anomalously thick) generally yield high calculated crystallization pressures (> 300 MPa = > 10 km), hydrothermal circulation (at least at slow-spreading ridges) must occur at pressures less than 300 MPa. More recent studies have now found Cl addition in basalts of some slow-spreading ridges (e.g. the Red Sea Rift) independent of their calculated crystallization pressures (Jenner and O’Neill 2012; van der Zwan et al. 2015) and advances in analytical methods mean that the inherently lower Cl contents of the relatively less fractionated slow-spreading ridge magmas have become less of an analytical challenge.

To investigate how deep hydrothermal circulation reaches at slower-spreading ridges, we performed high-precision Cl measurements (van der Zwan et al. 2012) on basaltic glasses and melt inclusions and combined this with barometric data based on glass compositions and on glass and clinopyroxene phenocryst compositions (cf. Ariskin and Barmina 2004; Putirka 2008). As the composition and structure of the new lithosphere at slower-spreading ridges is very heterogeneous, we selected two end-member ridges for our study: the magmatically robust southern Mid-Atlantic Ridge (SMAR) at 7–10°S (a slow-spreading ridge with ~ 3 cm year−1) previously interpreted to show varying mantle sources and crystallization pressures (0–800 MPa; Almeev et al. 2008; Hoernle et al. 2011) and the Arctic Gakkel Ridge at 6°W–85°E (an ultraslow-spreading ridge with max. 1.5 cm year−1) which has a strongly tectonic character and a very thin crust (Coakley and Cochran 1998; Dick et al. 2003; Michael et al. 2003). Note that at slow- and especially ultraslow spreading ridges classical “Penrose-type” magmatic crust (e.g. Vine and Moores 1972) may not make up the whole newly formed lithosphere (e.g. Cannat 1996) and may even be totally absent (e.g. Sauter et al. 2013). In this paper we therefore use the term “crust” to refer to the wholly magmatic upper part of the lithosphere, if present. The term “lithosphere” refers to the combination of crustal and/or mantle rocks which make up the new plate.

Geological background

Southern Mid-Atlantic Ridge

The southern Mid-Atlantic Ridge (SMAR) between 7° and 10°S, east of Ascension island, is a mature ridge with slow spreading rates of 31–32 mm year−1 full spreading rate (DeMets et al. 2010; Fig. 1a). The area is bounded to the north by the Ascension Fracture Zone and to the south by the Bode Verde Fracture Zone. The region is divided into four segments (A1–A4; Bruguier et al. 2003) that are separated by overlapping spreading centres and show varying morphologies. We chose to study samples from segments A1 and A2 as these samples display the largest range in calculated last crystallization pressures (200–900 MPa; Almeev et al. 2008). Segment A1 has a typical slow-spreading ridge character with a well-developed deep rift valley of > 2950 m, while segment A2 shows a shallower axial high and is 2100–3000 m deep (Fig. 1a; Bruguier et al. 2003). Segment A2 is associated with a thick seismically determined crust of ~ 11 km, compared to ~ 5 km under A1 (Minshull et al. 1998; Bruguier et al. 2003). Segment A1 displays microseismicity at 2–4 km depth and teleseismicity (Grevemeyer et al. 2013), while segment A2 shows little earthquake activity, thought to indicate a higher temperature gradient and therefore a shallower brittle/ductile transition in the crust (Fig. 1a; Devey et al. 2010). Presently, A2 appears to be in a period of reduced magmatic activity: the youngest lava flows are tectonized and yield ages around 1000 years and no signs of active hydrothermalism have been found (Devey et al. 2010; Haase et al. 2016). The samples used here are well studied in terms of their major and trace element and isotope compositions and crystallization pressures (Möller 2002; Almeev et al. 2008; Paulick et al. 2010; Hoernle et al. 2011). The geochemical studies show mixing between a depleted unradiogenic MORB source and radiogenic, trace element enriched HIMU-type sources (Hoernle et al. 2011). The enriched samples display higher H2O contents and nearly isobaric calculated crystallization pressures (100–300 MPa) while the more depleted samples show lower H2O contents and a range of calculated last crystallization pressures (200–900 MPa; Almeev et al. 2008). Hydrothermal circulation in the studied area is known from the Nibelungen hydrothermal field between A1 and A2 and from two indications of plume activity at 7°57′S and 8°17′S on A1 (Fig. 1a; Devey et al. 2005; Melchert et al. 2008).
Fig. 1

Overview of studied areas on the a Southern Mid-Atlantic Ridge (SMAR) and b Gakkel Ridge. The SMAR study area is subdivided in various segments (after Bruguier et al. 2003; samples are from A1 and A2). The Gakkel Ridge is segmented into three zones (after Michael et al. 2003): western volcanic zone (WVZ), sparsely magmatic zone (SMZ) and eastern volcanic zone (EVZ). White dots indicate earthquake hypocenters with M ≥ 3 (from ISC, 1960 to present). Black stars indicate hydrothermal activity after German et al. (2002) and Baker et al. (2004). Spreading rates (after DeMets et al. 2010) are indicated by arrows

Gakkel Ridge

The Gakkel Ridge (6°W–85°E) shows the lowest orthogonal spreading rate on Earth at 11–13 mm year−1 full spreading rate (DeMets et al. 2010) and lacks transform offsets (Fig. 1b). The Gakkel Ridge has a very deep rift valley (> 4600 m) and a thin seismically determined crustal thickness (< 4 km; Coakley and Cochran 1998; Jokat et al. 2003). The Gakkel Ridge can be subdivided into three different magmato-tectonic units (Fig. 1b; Michael et al. 2003). The Western Volcanic Zone (WVZ; 7°W–3°E) is a volcanic area with many small cones on the ridge and displays a larger extent of melting than the other two segments. In the Sparsely Magmatic Zone (SMZ; 3°E–28°E) in contrast, the seafloor consists mostly of peridotite with only rare basalts or volcanic features. The Eastern Volcanic Zone (EVZ; 29°E–85°E) contains large (~ 30 km) widely spaced volcanic centres of which some show recent activity. Basalts from the Gakkel Ridge are generally enriched in trace elements but with a large variability between the different segments and also between samples (Michael et al. 2003). Isotopically, two distinct sources were identified: samples from the WVZ have a DUPAL-signature attributed to a contribution from the sub-continental lithospheric mantle, while samples from the EVZ have a signature similar to Atlantic–Pacific MORBs (Goldstein et al. 2008) with the boundary between these mantle domains lying in the SMZ. Hydrothermal activity at Gakkel Ridge seems to be high compared to other (ultra)slow-spreading ridges (Edmonds et al. 2003; Michael et al. 2003). Although only two active vent sites were confirmed by visual observations (Aurora, 85°E), 7–11 more vent sites were inferred by light scattering, temperature and Mn concentration data in the water column, from which we here consider the seven most certain ones (Fig. 1b; Edmonds et al. 2003; Baker et al. 2004; Pontbriand et al. 2012). Teleseismic earthquakes are present along the whole ridge (Fig. 1b).

Analytical methods

Chlorine, major, trace, H2O and CO2 measurements and pressure calculations were performed on glassy basaltic samples from collections of expeditions R/V Meteor 41/2 (SMAR; e.g. Möller 2002; Almeev et al. 2008; Hoernle et al. 2011) and AMORE 2001 (Gakkel Ridge; e.g. Michael et al. 2003; Goldstein et al. 2008). For each basaltic glass measurement, three pieces of glass > 2 mm across were selected per rock under a binocular microscope based on their fresh appearance and lack of visible vesicles, alteration rims or minerals. The mounted glass was prepared, polished and then cleaned with Milli-Q water following the procedures in van der Zwan et al. (2012) to avoid Cl contamination. Optically clear SMAR Ol, Pl and Cpx minerals were prepared following the same procedure. For selected SMAR samples with elevated Cl/Nb glass contents, Ol and Pl minerals with melt inclusions were mounted in UHU Hart glue mixed with acetone and doubly polished sections (~ 60 µm thick) were prepared. Melt inclusions were selected based on homogeneity and large sizes/thicknesses and with a good surface.

Major elements and Cl measurements were carried out with a Jeol JXA-8200 ‘‘Superprobe’’ electron microprobe at GEOMAR, Kiel, using an acceleration voltage of 15 kV. Chlorine and potassium (K) were measured with a beam current of 80 nA and a beam diameter of 10 µm using the mapping technique described in van der Zwan et al. (2012, 2015). The Cl and K results are the average of 2–8 measurements per sample and have an uncertainty of < 3 and < 20 ppm for Cl and K, respectively (2SE). For the Cl and K measurements of the melt inclusions, the method was slightly adapted by a reduced beam current of 50 nA and a mapping repetition of 14 times, in order to avoid a strong Cl signal of the mounting material through the thin samples. Internal glass standards (138-3, 147-2, 157-3; van der Zwan et al. 2012) show a reproducibility of on average 3.2 ppm (2SD) for this method. Nevertheless, an additional signal from the mounting material is visible in the Cl data of the melt inclusions and corrected for, but results in a higher uncertainty of 19.2 ppm 2SD. This higher uncertainty does, however, not influence the trends observed in the melt inclusions, neither the main conclusions derived from Cl data of the melt inclusions. The K measurements are not affected by the mounting material. Any measurements that show discrepancies (e.g. inhomogeneous Cl, mostly due to bubble forming) were discarded.

Major elements of the glasses were measured with a defocused spot of 5 µm and a beam current of 10 nA. Counting times were 20/10 s (peak/background left and right of the peak each 10 s) for all elements apart for 30/15 s for MnO. Na2O was measured first to avoid loss by heating. For calibration and monitoring of data quality, we used natural reference samples from the Smithsonian Institute (Jarosewich et al. 1980). The average was taken of nine analyses per sample to assure homogeneity. Relative analytical precision is generally < 2.5%, but up to 5% for Na2O and ~ 30% for MnO and P2O5. Additionally, available (literature) data for the samples were used; for the SMAR, data are utilized from Möller (2002) and data measured with a CAMECA SX-50 at GEOMAR and a JEOL JXA8900 Superprobe at the Institut für Geowissenschaften, University of Kiel, partially presented in Haase et al. (2016). Settings were the same as above and accuracy and precision are better than 5% except for MnO and P2O3. Additional major element data measured on a Jeol JXA 8900RL electron probe microanalyzer at the University of Mainz were used for Gakkel Ridge basalts. Mineral point analyses and transects were measured with a focused beam spot. Plagioclase and Cpx were measured with a beam current of 20 nA and counting times of 20/10 s, apart for K2O in Pl (40/20 s) and Cr2O3 and TiO2 in Cpx (30/15 s). Olivine was measured with a beam current of 100 nA and 20/10 s counting times for SiO2, Al2O3, MgO, FeO, 30/15 s for MnO, Cr2O3, 60/30 s for CaO and 40/20 s for NiO.

For consistency, trace elements in all glass samples were analysed by LA-ICP-MS at GEOMAR, Kiel, using a 193 nm Excimer laser system (GeoLasPro, Coherent) coupled with a double-focusing, high-resolution magnetic sector mass spectrometer (AttoM, Nu-Instruments) under hot plasma conditions [NAI > 10; ThO/Th ~ 0.03%; details in Fietzke and Frische (2016)]. Spot analyses of samples were done by 30 s ablation at a laser repetition rate of 10 Hz using spot diameter of 90 µm and a fluence of 5 J cm−2. 50 s of gas background data were collected prior to each ablation. NISTSRM610 [30 s, 32 µm, 10 Hz, 5 J cm−2; Wise and Watters 2012] was used for calibration. MPI-DING reference samples [ML3B-G and KL2-G; 30 s, 32–90 µm, 10 Hz, 5 J cm−2; Jochum et al. (2006)] were used to evaluate the accuracy during the measurements. Relative analytical precision based on reproducibility (2SD) of the standard reference material used for calibration was about 3.5% (see Online Resource 1). Data evaluation has been performed applying the linear regression slope method (Fietzke et al. 2008). Ca and Si were used for internal standardization utilizing data from EMP analyses.

Glass H2O and CO2 concentrations were determined using a Fourier transformation infrared (FTIR) Bruker IFS88 spectrometer coupled with an IR Scope II microscope at the Institute of Mineralogy, Leibniz University of Hannover. H2O and CO2 concentrations were measured on doubly polished glass chips (~ 1 mm2 in size) that had a thickness of 150–50 μm, measured with a digital micrometer Mitutoyo [precision ± 2 μm (Behrens et al. 2009)] by three points for every polished glass fragment. Absorption spectra in the mid-infrared (MIR) range were collected using a spot size of 100 × 100 μm (average of 50 scans) and using the following operating conditions: globar light source, KBr beam splitter, MCT (HgCdTe) detector, 4 cm−1 spectral resolution, spectral range 13,000–0 cm−1. The H2O concentration was measured at the peak that is attributed to the OH stretch vibration (3550 cm−1) using a molar absorption coefficient of 67 L cm−1 mol−1 (Stolper 1982). The CO2 concentration was measured at the peak doublet with maxima at 1430 and 1510 cm−1 using molar absorption coefficient of 317 L cm−1 mol−1 (Shishkina et al. 2010). Glass densities were calculated using the equation of Yamashita et al. (1997) for basaltic compositions: ρ = 2819 − 20.8·H2O concentration. The H2O and CO2 concentrations were calculated based on the Lambert–Beer law using the peak height, which was determined by reference to a straight tangential base line. The average values of three measurements per sample were used to calculate the H2O and CO2 contents of the glasses, with a standard deviation usually less than 0.02 and 0.003 wt%, respectively.

Pressure and temperature conditions at the point of last crystallization equilibrium were calculated by two different methods. Pressures for all samples were calculated with the software COMAGMAT version 3.57 that is based on empirical models calibrated with an experimental database (Ariskin and Barmina 2004). Here we used inversed modelling (see Almeev et al. 2008) of the major elements and H2O concentrations of the glasses. The effects of small amounts of H2O on crystallization were taken into account by incorporating recent experimental calibration coefficients (Almeev et al. 2007a, b, 2012). The pressure models of COMAGMAT have an uncertainty of ± 100 MPa and are comparable with models used in Michael and Cornell (1998) (Almeev et al. 2008). For the samples that contain Cpx, pressures were also calculated using equilibrium exchange models between Cpx and the melt (Putirka et al. 1996; Putirka 2008), determined by analysing the average compositions of Cpx-rims and their host glasses. Mid-ocean ridge basalts have typically an Fe3+ content of 10–12% (Bézos and Humler 2005; Berry et al. 2017); for the calculations here we applied a Fe3+ content of 12% (Bézos and Humler 2005), but variations within these values do not affect the pressure calculations. The standard error of estimate (SEE) of the accuracy of the Cpx–melt model is ± 150 MPa (for an extended evaluation of involved errors see Putirka 2008).


Mineral textures and chemistry

Phenocrysts in the SMAR lavas are up to 2 mm in size and mostly comprise olivine (Ol) and plagioclase (Pl), with minor clinopyroxene (Cpx) observed in only five of the samples. In samples that cover all different chemical compositions Ol and Pl show a range of textures from euhedral crystals to anhedral grains displaying resorption at the mineral rims and occasionally sieve textures (Fig. 2a–e). A few of the Ol grains display undulose extinction (Fig. 2f). Melt inclusions are present in both Ol and Pl and can be up to ~ 200 µm in diameter; in the Pl they are frequently located along growth zones, but also along healed cracks. Olivine and Pl crystals often show clustering and intergrowth of multiple minerals that probably represent cumulate fragments (Fig. 2c–h).
Fig. 2

Photomicrographs of thin sections from SMAR samples 149-1 (a, cf, h), 142-7 (b) and 145-3 (g) in cross-polarized light (ad, f–h) and plane-polarized light (e). a, b, h Euhedral Ol occurs together with highly dissolved and skeletal Pl and/or Ol. ch Clusters of multiple Pl (c), or Pl and Ol (dh) with complex intergrowth structure of Pl and Ol (e, g) and undulose extinction in Ol (f) indicate xenocrystic fragments that formed by in situ crustal growth of minerals together and subsequent assimilation of these fragments

The forsterite (Fo) contents in Ol vary between 83.4 and 91.8%, with intra-sample variations up to 5.5% both between different grains and within single Ol grains. Chemical profiles from core to rim over several selected Ol reveal (1) crystals with a constant composition, (2) normally zoned crystals, of which some may include diffusion profiles (e.g. in 143-2 and 149-1) and (3) crystals with a slightly reverse zonation (e.g. in 143-2, 147-2; Fig. 3a). Several Ol rim compositions are too Fo-rich to be in equilibrium with their host melts (Fig. 3b; using the equilibrium constant of 0.30 ± 0.03 from Roeder and Emslie (1970) and 12% Fe3+ in the glasses (Bézos and Humler 2005) points to equilibrium at Fo-contents of 79.4–89.6%) implying a xenocrystic origin for these Ol. Also here applying a slightly different Fe3+ content does not affect this observation. Olivines in disequilibrium are found for both incompatible element enriched and depleted samples. The rare Cpx that occur in samples 147-2, 149-1, 149-4, 162-1 and 162-2 have a composition of Wo36–41En51–55Fs6–10 and rim to core profiles display zoning in some cases (Fig. 3c). Two profiles over Pl show oscillatory zoning around a constant composition or slight reversed-zoning (Fig. 3d).
Fig. 3

Mineral composition data. a Transects over Ol grains from rim to core show unzoned Ol together with normally and slightly reversely zoned Ol. b Fo-content of the Ol plotted against the Fo-content of Ol that would be in equilibrium with their host glasses (applying the equilibrium constant of 0.3 ± 0.03 from Roeder and Emslie 1970). Olivines in equilibrium fall onto the one-to-one line or within the red field that represents the variation of the equilibrium constant. c Transects over Cpx crystals show various zonation patterns from limited zoned to complexly zoned crystals. d Plagioclase transects display scatter showing oscillatory zoning, but display a generally reversely zoned trend (147-2) and unzoned trend (143-2)

Composition of the glasses

The major and trace elements together with the Cl data are given in Online Resources 2 (SMAR) and 3 (Gakkel Ridge). The Cl concentration values range from 43 to 704 ppm for Gakkel Ridge and 52–429 ppm for the SMAR (± < 3 ppm 2S.E.), although most values lie below 200 ppm (Fig. 4a). This range is similar to literature data for other slow-spreading ridges (Michael and Cornell 1998; PetDB, Lehnert et al. 2000; Jenner and O’Neill 2012). In the majority of both SMAR and Gakkel samples there is a negative correlation between Cl and MgO (Fig. 4a). This trend is, however, poor (R 2 = 0.22; n = 117) and not parallel to a single fractionation model [(log(Cl8/Cl) = 0.12·(MgO-8) from Weaver and Langmuir (1990), assuming unfractionated Cl concentrations of 50, 100 or 150 ppm at 8 wt% MgO; Fig. 4a]. Several Cl-rich samples from both SMAR and Gakkel Ridge lie well above the values that could be explained by fractionation. Positive correlations are seen between Cl and several incompatible elements, e.g. K, Rb, Ba, Th, U, Nb, Ta, La, Ce, Pr, Nd and H2O, with the best correlation existing between Cl and Nb (R 2 = 0.87; Fig. 4b, c), the element whose mantle incompatibility is most similar to Cl (Sun et al. 2007). Samples from the SMAR also show a correlation between Cl and radiogenic Sr, Nd and Pb isotopes (isotope data from Hoernle et al. (2011); Fig. 4d; Table 1). No correlation (R 2 < 0.2) is observed between Cl and indicators of degassing (e.g. CO2/Nb, CO2/Ba, cf. Saal et al. 2002; Michael and Graham 2015), melting degrees (e.g. Na8, cf. Klein and Langmuir 1987) or continental input (e.g. Ce/Pb, U/Nb, cf. Hofmann et al. 1986). A good correlation between H2O and Ce (Fig. 4e) suggests that the magmas have not lost or gained significant amounts of water since leaving the melting zone (cf. Michael 1995).
Fig. 4

Chlorine compositions of SMAR and Gakkel Ridge basalt glasses compared to major and trace element compositions. The analytical error is smaller than the symbol sizes. a Cl vs. MgO compared to fractionation trends for Cl compositions of 50, 100 or 150 ppm at 8 wt% MgO (black lines). Cl vs. Nb for b Gakkel Ridge and c SMAR samples together with lines of various Cl/Nb ratios and compared to samples from the Red Sea (van der Zwan et al. 2015). The local mantle Cl/Nb ratio (see “Discussion”) for the two locations is highlighted. Note that many samples lie above these mantle values. d Chlorine concentrations are related to radiogenic isotope ratios (i.e. 87Sr/86Sr; Hoernle et al. 2011). e H2O and Ce contents are relatively well correlated. f Cl/K and Cl/Nb are correlated for each sample suite, but display different trends between the suites. g Variations in Nb/K between the samples suites are corresponding to Ce/Pb

Table 1

Pressure calculations based on the formulations for Cpx–melt equilibria compared to COMAGMAT pressure calculations



Pressure Putirka et al. (1996) (MPa)

K D (Fe–Mg)

Pressure COMAGMAT (MPa)

∂ Pressure






− 309






− 354






− 346



− 13



− 18











































aThe low/negative pressures together with the low K D (Fe–Mg) indicate disequilibrium and geologically non-significant pressures for this sample

To study Cl addition to the magmas, we examined the ratios of Cl to elements of similar incompatibility; wherever possible we used Nb (Cl/Nb) as Nb is slightly closer in incompatibility to Cl than K (Sun et al. 2007) and is more resistant to the effects of weathering and alteration processes; for analytical reasons Cl/K is used for the melt inclusions, and to compare our data to older studies, where mostly only major elements were measured. Cl/K and Cl/Nb are positively correlated for each study area although each area lies on a different subparallel trend (Fig. 4f). These trends are consistent with source-derived variations in Nb/K between the different sample suites as they also show a grouping in their Ce/Pb and Nb/U contents (Fig. 4g). SMAR samples reach a maximum Cl/K of 0.3 and Cl/Nb of 64, higher than at most other slow-spreading ridges, but lower than has been observed at fast-spreading ridges or in the Red Sea (Fig. 5). The Gakkel Ridge samples, in contrast, all lie within the previously defined slow-spreading ridge field (Michael and Cornell 1998; Fig. 5). High Cl/K ratios at SMAR mostly occur in trace element depleted samples (e.g. samples with low K/Ti, Fig. 5a), consistent with the negative relation observed between Cl/Nb and incompatible trace elements (e.g. Nb, K, Rb, REE; Fig. 5c). Samples from Gakkel Ridge show a range of Cl/Nb values across the whole range of trace element concentrations (Fig. 5b). High Cl/Nb is observed in both segments of the SMAR and in the WVZ and SMZ of the Gakkel Ridge. The EVZ of the Gakkel Ridge in contrast, displays comparatively few high Cl/Nb samples.
Fig. 5

Indicators of hydrothermal Cl addition (Cl/Nb, Cl/K) in basaltic glass plotted against major and trace element data. The analytical error is smaller than the symbol sizes. a Cl/K ratios reach higher values than previously observed for average slow-spreading ridges, but are not as high as observed for Red Sea Rift basalts or fast-spreading ridges. High Cl/K (> 0.1) is not observed for samples with K/Ti > 0.22. Numerous b Gakkel Ridge and c SMAR samples have Cl/Nb above the Cl/Nb mantle ratio. Samples from Gakkel Ridge show high Cl/Nb over the complete range of Nb; at the SMAR this is mainly for samples with low Nb

High-precision chlorine measurements on SMAR melt inclusions

High-precision Cl and K concentrations were measured in melt inclusions in both Ol and Pl crystals in selected high Cl/Nb (28–48) samples from the SMAR (137-3, 143-1, 143-2, 147-2, 147-3, 149-1, 149-4 and 151-3; Online Resource 4). The SMAR melt inclusions display a generally positive trend between Cl concentrations (20–168 ppm in Ol, 30–234 ppm in Pl) and K (67–2256 ppm), but with more scatter towards higher Cl-values than observed in the bulk glasses (Fig. 6a). This scatter is not concentration-dependent, and is considerably larger than analytical error. Cl/K ratios in the inclusions (0.08–0.56) extend mainly to higher values than observed in the host glasses (0.12–0.19). For individual samples, Cl, K and Cl/K of the melt inclusions vary to both higher and lower values than their host glass (stars in Fig. 6b). An exception are the trace element enriched samples from station 149, which display a larger variation of the melt inclusions in Cl/K, but both Cl and K are considerably lower than the host glass and overlap with the more depleted samples (Fig. 6c). The highest Cl/K values in the melt inclusions are observed in 143-1 and 143-2 that also have the highest Cl/K of their host glasses (0.19).
Fig. 6

Melt inclusion Cl and K data. a Melt inclusions in Ol (diamonds) and Pl (circles) show a similar trend in Cl and K as the SMAR glasses, but with larger scatter. There is no systematic difference between melt inclusions in Ol or Pl. b Cl/K vs. K for melt inclusions (filled symbols) together with their host glasses (stars; colours distinguish samples). Melt inclusions in minerals hosted in trace element depleted glasses display both lower and higher Cl/K and K than their host glasses. Note that the most K-depleted inclusions show extremely high Cl/K ratios, exceeding the range in their host glasses. c Melt inclusions from trace element-enriched (high K) glass sample 149-1 have lower Cl and K than their host and overlap with concentrations of melt inclusions in trace element depleted samples (b)

Pressures of equilibrium crystallization

Previous work (Almeev et al. 2008) had used the experimentally calibrated COMAGMAT program to determine the pressures at which the SMAR magmas were last in equilibrium with the three-phase assemblage Ol–Pl–Cpx. Applying COMAGMAT with improved experimental calibration coefficients for the effect of water (see method section) to our samples yields pressures of last equilibrium crystallization between 5 and 700 MPa for the SMAR and between 100 and 840 MPa for the Gakkel Ridge (Online Resources 2, 3). The average pressures are 357 and 426 MPa, respectively, similar to values found by Michael and Cornell (1998) for slow-spreading ridges which did not display high Cl/K. SMAR segment A1 has slightly higher calculated crystallization pressures than the shallower segment A2 with on average 385 and 322 MPa, respectively. At Gakkel Ridge, the highest average pressures are found in the Eastern Volcanic Zone (538 MPa), while the Western Volcanic Zone and Sparsely Magmatic Zone display lower pressures (340 and 375 MPa, respectively).

For the rare samples that contain Cpx (SMAR 147-2, 149-1, 149-4, 162-1, 162-2), we also calculated pressures (given in Table 1) using the formulations for Cpx–melt equilibrium of Putirka et al. (1996). Melt–phenocryst equilibrium was assessed by comparing the predicted and observed values for the Cpx components (cf. Putirka 2008) and by comparing the observed and predicted values for K D (Fe–Mg)cpx–liq (Online Resource 5; Table 1). In terms of the Cpx components, the calculated and observed compositions were within 5% and indicate equilibrium in all samples. Samples 147-2 and 162-2 also show equilibrium in K D (Fe–Mg)cpx–liq. Samples 149-4 and 162-1 display minor [K D(Fe–Mg)cpx–liq = 0.17–0.20] and sample 149-1 major [K D(Fe–Mg)cpx–liq = 0.12] disequilibrium in Fe and Mg of the Cpx with the glass, indicating some late-stage modification of the glass for these fast-diffusing elements. As the pressure calculations are dependent on variations in the Cpx components, which involve slower diffusing elements, this minor disequilibrium in Fe–Mg for samples 149-4 and 162-1 has no significant effect on the resulting pressure calculations. The disequilibrium effect is significant, however, for sample 149-1 so we have excluded it from the subsequent discussion. The four samples in (close) equilibrium yield pressures between 300 and 400 MPa, which are not consistent with the pressures calculated by COMAGMAT (Table 1). This will be discussed further below.


Chlorine contents and Cl/Nb values of the mantle and magma

In order to identify Cl addition to the magmas from external processes, it is necessary to determine the possible range in local mantle Cl contents (and hence its probable concentrations in mantle-derived magmas). The correlation between Cl and many incompatible trace elements (particularly those that have a compatibility close to Cl, for example Nb) and to Sr-, Nd- and Pb-radiogenic isotopes in the case of the SMAR (Fig. 4b–d) indicates that much of the variation in absolute Cl concentrations between the samples can be explained by source processes. Mantle source variations at the SMAR were described by, e.g. Almeev et al. (2008) and Hoernle et al. (2011) and were attributed to mixing between a depleted mantle source (with low Cl, Nb and K concentrations) and several enriched (HIMU) mantle sources (with higher Cl, Nb and K concentrations), while variations in mantle source compositions at the Gakkel Ridge were identified by, e.g. Michael et al. (2003) and Goldstein et al. (2008). We see from Fig. 4f, g that each individual axial locality has its own inherent Nb/K signature and, by inference, Cl/Nb and Cl/K mantle signatures. If we can define these mantle ratios for the individual regions, then Cl/Nb and/or Cl/K above these values should indicate Cl addition.

Cl/Nb (and Cl/K) mantle ratios are not well constrained and cover a large range in the literature: Palme and O’Neill (2003) suggest the highest Cl/Nb of 51, McDonough and Sun (1995) propose 25.8, le Roux et al. (2006) favour 14, Salters and Stracke (2004) choose 6.9 and Saal et al. (2002) give a ratio of 3. However, Palme and O’Neill (2003) calculate a relatively high Cl mantle value (30 ppm) compared to values that were derived from basaltic melts and melt inclusions, which are all lower than 7 ppm (Ryabchikov 2001; Saal et al. 2002; Salters and Stracke 2004; Kovalenko et al. 2006; le Roux et al. 2006), while values derived from mantle minerals are even lower < 0.38 ppm (Urann et al. 2017). If we assume mantle concentrations of 0.3–0.6 ppm for Nb (Hofmann 1988; McDonough and Sun 1995; Palme and O’Neill 2003) and a Cl concentration of < 7 ppm, then mantle Cl/Nb ratios should lie at least below the value of 25.8 proposed by McDonough and Sun (1995). To determine in more detail the local Cl/Nb in our study areas we can use the variations we see in our samples. Assuming (a) no fractionation of Cl/Nb during melting (which can be excluded due to their similar distribution coefficients) and (b) no degassing of Cl from the magma prior to eruption (see below) then interactions of the magma with its environment should always lead to Cl/Nb increases. In this case, the magmas with the lowest Cl/Nb represent the least contaminated samples and their Cl/Nb should be closest to the local (upper limit) mantle Cl/Nb ratio. The evidence for a mix of sources at both locations may imply that the Cl/Nb and Cl/K mantle ratios are not constant within an area, which may be particularly the case for the SMAR ridge samples as HIMU sources may have higher Cl/K values (Stroncik and Haase 2004; Kendrick et al. 2017). However, Fig. 5b, c, shows a very consistent lower Cl/Nb ratio in our samples over a large range of absolute Nb contents at both locations. At Gakkel Ridge these values span the full range of compositions and a constant Cl/Nb mantle value is confirmed by the fact that we see, for example, no correlations between Cl/Nb and any trace element ratios indicating variations in continental input (e.g. Ce/Pb, U/Nb, cf. Hofmann et al. 1986) at the Gakkel Ridge.

In the SMAR samples we see a negative correlation between the lowest Cl/Nb values and trace element concentrations as all depleted SMAR basalts (Nb < 4.5) show higher Cl/Nb ratios (significantly above analytical error, see “Analytical methods”), while SMAR samples with low Cl/Nb and Nb > 4.5 ppm fall on a line with Cl/Nb of 20 ± 2 (Fig. 5b). We cannot rule out that this results from mixing of a magma from a low Nb—high Cl/Nb (~ 50) source with magma from a trace element enriched, low Cl/Nb source, but this appears unlikely for the following reasons: (a) we see a range of Cl/Nb values for similar trace element (Nb) concentrations that cannot be explained by mantle variations; (b) we see no positive correlation between Cl/Nb and other indicators of mantle enrichment (e.g. Zr/Hf; Nb/Zr); (c) depleted mantle sources have been shown to have lower Cl/K (and by implication, Cl/Nb) than enriched, particularly HIMU, mantle sources (Michael and Cornell 1998; Stroncik and Haase 2004; Kendrick et al. 2017). Therefore, if the depleted mantle source (low Nb samples) has a different Cl/Nb ratio, this is expected to be lower than the Cl/Nb of 20 ± 2 indicated by the HIMU source samples (high Nb values). Instead, the higher Cl/Nb for all SMAR samples with Nb < 4.5 ppm, more likely can be explained by their inherently higher sensitivity to Cl contamination due to their low magmatic Cl (the effect of adding a constant amount of excess Cl to magmas with different Nb concentrations are shown as dashed lines on Fig. 5b, c; note that the smoothness of the curves observed are an artefact of plotting the ratio of an element against the same element). We thus take these values to indicate the local mantle Cl/Nb ratios (ca. 20 ± 2 at SMAR, 18.5 ± 1.5 at Gakkel). All samples with Cl/Nb above these local mantle Cl/Nb ratios are taken to show Cl-excess (defined as Cl content above the expected Cl in the melts based on their Nb content and the local Cl/Nb mantle ratio).

Using the correlation regression shown in Fig. 4f, Cl/Nb mantle ratios of 18.5 ± 1.5 and 20 ± 2 for the Gakkel Ridge and SMAR, respectively, correspond to Cl/K mantle ratios of 0.040 and 0.086 (Fig. 6b, c), which is consistent with typical Cl/K values of 0.02–0.04 and 0.04–0.08 determined by Stroncik and Haase (2004) for EM (Gakkel Ridge) and HIMU (SMAR) suites.

Using these Cl/Nb mantle ratios, we can calculate Cl-excess in ppm for each sample based on its Nb or K content (Fig. 7a, b). In contrast to the absolute Cl/Nb ratios, Cl-excess is independent of the trace element (e.g. Nb) concentrations of the samples, and shows only the addition of Cl to magmas derived from any particular mantle source. Cl-excess is up to 250 ppm, and ~ 75% of the measured samples from both the Gakkel Ridge and SMAR show Cl addition (Online Resources 2, 3) over a range of compositions, spreading rates and both comparatively thick and thin crust (Fig. 7). The amount of Cl-excess is similar for the SMAR and Gakkel Ridge, apart from samples from the EVZ that show significantly less Cl-excess (Fig. 7a). The SMAR melt inclusions show the same extent of Cl-excess as the glasses (Fig. 7b). Both ridges show considerably lower Cl-excess than basalts from the slow-spreading Red Sea (< 1400 ppm; Fig. 7a; cf. van der Zwan et al. 2015).
Fig. 7

Cl-excess (ppm) calculated by the Cl concentrations minus the magmatic Cl based on Cl/Nb mantle values plotted against major and trace element data. a Cl-excess is present in the glasses over various Nb concentrations and is up to 250 ppm. b The melt inclusions show the same of Cl-excess as the glasses. c Cl-excess is not related to MgO (fractionation), d Na8 (melt degree), e 87Sr/86Sr (source variation) or f CO2 (degassing)

Chlorine excess and magmatic, seafloor and hydrothermal processes at SMAR and Gakkel Ridge

Magmatic processes such as fractional crystallization and partial melting as well as source variations influence the absolute Cl contents of the melts, but do not affect the Cl/Nb ratios and hence cannot cause the observed Cl-excess. This is additionally confirmed by the lack of a correlation between Cl-excess and indicators of fractionation (MgO) or degree of melting (Na8) and source variations (87Sr/86Sr; Fig. 7c–e). Degassing of the samples is indicated by the CO2/Nb of the samples of mostly < 200 and the CO2/Ba of < 60 that are significantly lower than values of 239 or 283 for CO2/Nb and 105 for CO2/Ba for undegassed magma derived by Saal et al. (2002) and Michael and Graham (2015). Degassing of CO2 can also be seen in a plot of CO2 against H2O compared to isobars of equilibration eruption pressures where we see degassing trends for CO2 towards the eruption pressures of the samples (Online Resource 6). H2O shows, however, no degassing trends and is undersaturated compared to the isobars, implying no H2O degassing took place, consistent with the good H2O/Ce correlation. Also variations in Cl due to Cl removal by degassing can be ruled out as there is no correlation between CO2/Nb or CO2/Ba and Cl-excess (Fig. 7f).

As melt inclusions display the same Cl-excess as the glasses (Figs. 5a, 6, 7a, b) and Cl-excess is homogeneous across multiple glass chips of the same sample and also not correlated to indicators of surficial alteration e.g. U, Ba (e.g. Alt et al. 1986; Alt and Teagle 2003; Schramm et al. 2005; Augustin et al. 2008), Cl addition by surficial processes (seafloor alteration or syn-eruptive magma–seawater interaction, e.g. Hart et al. 1974; Soule et al. 2006) can be ruled out as playing a significant role in generating the Cl-excess we observe. Instead, the source of the Cl-enrichment must lie deeper in the magmatic system, before eruption but during crystal growth and magma homogenization.

The assimilation of hydrothermally altered oceanic crust was shown to lead to Cl-excess in glasses at other locations (cf. Bédard 1991; Bédard and Hébert 1996). Furthermore, the global lack of Cpx in most MORB magma despite a positive correlation between Ca and Mg has been explained by assimilation of (lower) crustal rocks (e.g. Kvassnes and Grove 2008; Lissenberg and Dick 2008; Godard et al. 2009; Drouin et al. 2009) and melt–rock reactions (syntexis; Bédard et al. 2000 and references therein; Kelemen et al. 1995). The petrology of our SMAR samples shows that they have a complex history of melt/rock interactions, with phenocryst rim/glass disequilibrium (Fig. 3b), complex or reverse zoning and dissolution features (Fig. 2) all suggesting magma mixing and xenocryst inheritance. The presence of depleted melt inclusions in phenocrysts from enriched magmas (Fig. 6c) further underlines their xenocrystic origin and the late addition of the enriched component to the system before eruption (Fig. 6c). The occurrence of rock fragments identified by deformation textures in the SMAR samples (Figs. 2, 3) suggests that the magmas assimilated oceanic crust or partially solidified crystal mushes. Evidence for assimilation of country rocks at Gakkel comes from U-series data (Elkins et al. 2014).

To determine whether the Cl-excess in our glasses comes from assimilation of country rock during ascent we can use the H2O/Cl ratios of the samples, as this ratio is different in seawater, altered rock and hydrothermal brines. When we plot H2O/Cl against K/Cl (following Kendrick et al. 2013; Fig. 8), we see that our samples follow a mixing line from mantle values to a contaminant with low H2O/Cl (< 9). Both hydrothermally altered lithosphere and seawater have high H2O/Cl ratios (> 16 and ~ 50; see Ito et al. 1983; Bach et al. 2003; Sano et al. 2008; Barnes and Cisneros 2012) and would introduce H2O to the magma. The good correlation between H2O and Ce (maximum 0.22 wt% variation for a given Ce content per sample suite; Fig. 4e), the undersaturation of the magmas in H2O (Online Resource 6), and H2O/Ce values of < 350, all indicate no significant H2O addition (see also Michael 1995; le Roux et al. 2006; Wanless et al. 2010, 2011; van der Zwan et al. 2015). The Cl-excess we observe cannot, therefore, be explained by the bulk addition of either seawater or bulk altered lithosphere. A Cl-rich phase (that adds up to 250 ppm Cl) with low H2O/Cl can be introduced by assimilation of 0.25–0.08% fluids with NaCl contents of 16 to > 50%, respectively (brines, e.g. Michael and Schilling 1989; Kendrick et al. 2013; van der Zwan et al. 2015; Fig. 8). Such fluids are generated, particularly at high temperatures and pressures, by phase separation of hydrothermal fluids (Bischoff and Rosenbauer 1987; Fournier 1987; Berndt and Seyfried 1990). Hydrothermal brines could potentially be added directly to the magmas if fluids intrude through the conductive cracking front of the magmas (e.g. Kendrick et al. 2013); the petrographic evidence for assimilation of country rocks in both sample sets argues, however, against such a simple incorporation of fluids. Assimilation of hydrothermally altered rocks with brines present in the pore-spaces (Coogan et al. 2003) or as fluid inclusions in hydrothermally formed minerals, would lead to Cl-excess but little H2O addition. This mechanism for Cl-contamination has also been suggested for samples at the boundary between the magmatic and hydrothermal systems from IODP Site U1256 (fast-spreading EPR), where Cl was found to be present in fluid inclusions or pore fluids (Zhang et al. 2017) and stoping of altered crust was thought to be the dominant Cl enrichment process (Fischer et al. 2016). Alternatively, Cl could be added by partial melting (anatexis) of the hydrothermally altered lithosphere and particularly by the breakdown of Cl-rich hydrothermally formed minerals such as amphiboles (Barnes and Cisneros 2012); these minerals melt preferentially due to a lower solidus than the fresh rock, and partition Cl much more strongly than H2O into the melt, producing a low H2O/Cl contaminant (Michael and Schilling 1989; Kent et al. 1999; France et al. 2010, 2014; Wanless et al. 2010, 2011). This process has been proposed for some basaltic to dacite lavas on the EPR based on their major and trace element data (France et al. 2010, 2014; Wanless et al. 2010, 2011) although it was excluded for the rocks from IODP Site U1256 which apparently reacted under dry conditions without the influence of amphibole (Fischer et al. 2016; Erdmann et al. 2017). Since there is no correlation in our samples between Cl-excess and any of the other measured elements (Fig. 7; Online Resources 2, 3) that could be indicative for any of these processes, we cannot distinguish, with our data, between the assimilation of altered country rock containing interstitial brines or brine in fluid inclusions and partial melt of (potentially amphibole-bearing) altered lithosphere. We therefore group all these potential deep Cl contamination processes under the term ‘assimilation of hydrothermally altered lithosphere’.
Fig. 8

The relation of the K/Cl and the H2O/Cl contents of the samples compared to their mantle values and potential contaminants (seawater, altered oceanic crust, brines). Salinity of the brines with 10–50 wt% salts and mantle data are after Kendrick et al. (2017), but with the local K/Cl mantle values as derived in this paper. Altered oceanic crust data are from Ito et al. (1983), Bach et al. (2003), Sano et al. (2008), Barnes and Cisneros (2012). See text for discussion

Crystallization depths and the validity of the pressure calculations

The discrepancies between the pressure of last equilibrium calculated using COMAGMAT (0–840 MPa) and those derived from the Putirka et al. (1996) Cpx–melt equilibrium model (300–400 MPa) indicate that at least one type of pressure estimate is inaccurate. The COMAGMAT pressures are representative for melt-based pressure calculations (see Online Resource 7 for discussion), but the assumption, inherent in all phase-based pressure calculations, of a simple crystallization history and equilibrium at the Ol–Pl–Cpx cotectic as the defining final event in the magmatic evolution is probably incorrect. Petrographically and geochemically our samples show signs of mineral–melt chemical disequilibrium (Fig. 3), assimilation of lithic xenoliths (Fig. 2) and magma mixing (Figs. 3, 6). Moreover, the lack of Cpx phenocrysts in most of the samples is inconsistent with the assumption that the melts are at a multiply-saturated cotectic with Ol, Pl and Cpx.

Melt–rock reactions in the lower lithosphere may significantly change the compositions of both the cumulates (e.g. Meyer et al. 1989; Bédard 1991; Bédard and Hébert 1996; Bédard et al. 2000; Coogan et al. 2000; Dick et al. 2002; Ridley et al. 2006; Grimes et al. 2008) and the melt. As we have no precise information on the original melt and contaminant compositions, the extent of assimilation cannot be determined straightforwardly by backward-modelling although Kvassnes and Grove (2008) and Lissenberg and Dick (2008) have shown, using experiments and forward modelling, that assimilation of various types of oceanic crust has a significant effect on the composition of the melts and their Al2O3–CaO–Si–MgO relations. They concluded that any method that uses the composition of a melt (glass) alone to determine, e.g. crystallization pressure or temperature from a melt can be erroneous. Lissenberg and Dick (2008) further showed that calculated pressures increase with increasing assimilation of lower crustal troctolite (up to 25%) and that the full range of crystallization pressures observed in MORB (100–800 MPa) can be achieved by reaction of melts with the lower crust. This is consistent with the basalt data from the SMAR, where many of the samples that display high COMAGMAT pressures (> 300 MPa; Table 1) show Cl-excess implying they have assimilated altered crust. As all samples with Cl-excess likely have modified major element compositions due to assimilation, pressure calculations based on major elements cannot be used to determine the depth of Cl addition.

Pressure calculations based on Cpx–melt element exchange (Putirka 2008) can give a better control on the crystallization pressures, as the degree of equilibrium between glass and Cpx can be assessed and, by using the Cpx rim compositions, only equilibrium in the last phase of crystallization is required, thus the pressures are not affected by earlier magma modifications. A statistically robust assessment of the relationship between Cl-excess and crystallization pressures is hindered by the lack of Cpx in most (MORB) samples, but our samples with Cpx in equilibrium give geologically relevant depths.

The depth of assimilation of hydrothermally altered lithosphere

The SMAR Cpx–melt pressures (after Putirka et al. 1996) all lie in a relatively restricted range (316–403 ± 150 MPa ≈ 10–13 ± 5 km; Table 1). For a crustal thickness of ~ 11 km on segment A2 (Bruguier et al. 2003) where these samples originate, these pressures imply that the last episode of crystallization occurred at lower crustal depths or close to the Moho. The Cl-enriched melt inclusions in Ol and Pl imply an increase of Cl in the magma prior to crystallization and inclusion formation. The complete equilibrium observed for samples 147-2 and 162-2 between the Cpx phenocrysts and melt show no substantial modification of the melt subsequent to Cpx crystallization. Thus Cpx crystallized together or after the Ol and Pl that host the Cl-contaminated melt inclusions. This places a strong limit on the minimum depth of hydrothermal circulation in this region of ~ 10 km. We note that not all samples that display lower crustal Cpx crystallization pressures display Cl-excess, indicating that not all assimilated crust is altered or that assimilation of crust is not the only method of magma cooling at depth.

Also the investigated samples that show a slight modification of the host melt after crystallization of the Cpx of the faster diffusing elements by minor disequilibrium in Fe–Mg between the Cpx and glass, but with Cpx components that are in equilibrium, host melt inclusions with Cl-excess indicating deep assimilation of brines (Table 1). Disequilibria between melts and several Ol rims imply assimilation or mixing processes shortly before eruption, as most Ol in disequilibrium do not show any evidence of diffusive re-equilibration that would have led to the formation of obvious diffusion profiles within a few days at magmatic temperatures (cf. Klügel 1998; Shaw and Klügel 2002). Even the one observed diffusion profile with a diffusion length of < 50 µm (sample 143-2; Fig. 3a), indicates a short residence time of the Ol in the melt of maximum a few weeks (cf. Shaw and Klügel 2002). The lack of diffusion profiles in the Ol (Rutherford 2008) shows that the rising of magma was relatively fast (days to a few weeks) and therefore Cl addition by assimilation of crust is less likely to happen during ascent in the upper crust. This is in agreement with what would be expected at slow-spreading ridges, where the cooler upper crust (Reid and Jackson 1981; Dick et al. 2003; Montési and Behn 2007) requires that magma rises fast in order to not crystallize.

Clinopyroxene phenocrysts were only found at SMAR A2, where the anomalously thick crust probably favours the early crystallization of clinopyroxene. At slow-spreading ridges with typical or lower crustal thickness without clinopyroxene phenocryst (SMAR segment A1 and Gakkel Ridge) crystallization and assimilation depths cannot be directly determined. Nevertheless, the ranges in Cl, K, Cl/K and Cl-excess observed in melt inclusions in individual samples also from SMAR segment A1 imply that they were formed from both different mantle melts and melts which had experienced varying degrees of assimilation before extensive crystallization (Figs. 6, 7b). As many Ol grains are xenocrystic, they must have crystallized from earlier batches of melt that already assimilated Cl as shown by the Cl-excess of the melt inclusions in the Ol. Hence, the various amounts of Cl-excess observed in the glass and melt inclusions from one rock probably represent multiple generations of hydrothermal alteration, assimilation of crust and crystallization, rather than processes occurring all to the same magma batch. Production of the crust by multiple cycles of crystallization was also suggested by Grimes et al. (2008) and Lissenberg et al. (2009), who found a range of ages in gabbros from the Mid-Atlantic Ridge.

We see very similar degrees of Cl-excess in all regions studied and the homogeneous distribution of Cl in the bulk-rock glasses from both the SMAR and Gakkel Ridge shows extensive homogenization after assimilation, which requires magma reservoir processes. Therefore, also at Gakkel Ridge, Cl cannot simply have been added during ascent, implying that hydrothermal circulation must reach the levels of magma pockets at depth. At the ultraslow-spreading Gakkel Ridge, the thin lithosphere (Reid and Jackson 1981; Dick et al. 2003; Montési and Behn 2007) will be inherently more strongly cooled by conduction, which promotes deep crystallization and so magma reservoirs are expected to form deeper there (Rubin and Sinton 2007; Morgan and Chen 1993). Both deep pooling of melt (Shaw et al. 2010) as well as melt–rock interactions at depths < 15 km (Elkins et al. 2014) have been suggested at Gakkel Ridge. This evidence and the consistent level of Cl-excess between the ridges implies that deep hydrothermal circulation is also occurring there. Crystallization and crustal assimilation in magma bodies in the lower lithosphere at (ultra)slow spreading ridges is consistent with the evidence of gabbroic samples and melt–rock reactions in the lower crust and mantle found at the South-West Indian Ridge and the MAR 30°N (e.g. Dick et al. 2000; Expedition Scientific Party 2005).

Deep circulation of water can be facilitated by deep faults-pathways. Data for teleseismic earthquakes (1960 to present; International Seismological Centre 2011) record events at depths of > 10 km for both study areas. Although the depth estimate of these earthquakes is not very precise (Engdahl et al. 1998), the > 3 magnitude of these earthquakes (Fig. 1) indicates that the faults causing them must be large and thus have the potential to extent to larger depths. Also, microseismicity on SMAR segment A1 down to 2–4 km within the 5-km-thick crust, indicates that faults extend to lower crustal depths (Grevemeyer et al. 2013). Deep faults, down to the lower crust, were found in gabbros at slow-spreading ridges, e.g. by Harper (1985) Cannat et al. (1991) and Mével and Cannat (1991), while Bach et al. (2004) showed that also serpentinites that may be present in the lithosphere of ultraslow-spreading ridges can break and alter at depth. In addition, Hasenclaver et al. (2014) showed that the physical process of deep hydrothermal circulation can be modelled at greater depths.

In conclusion, hydrothermal circulation is not restricted to the upper crust and the assimilation depth is therefore not a limiting factor on Cl-addition to rising magmas as suggested by Michael and Cornell (1998). The difference between our conclusion and those of previous work is probably related to analytical precision, which needs to be high to distinguish between small, but significant variations in Cl/Nb and to identify appropriate Cl/Nb mantle values and Cl-excess in slow-spreading-ridge magmas.

Chlorine excess in lavas and sites of hydrothermal activity

To study the spatial relationship between Cl-excess in lavas and hydrothermal vent sites, we compared the sampling locations of basalts showing Cl-excess with those of known active hydrothermal vent sites and indirect indications for hydrothermal activity (e.g. plume signatures; Fig. 9). At both the SMAR and Gakkel Ridge there are distinct areas with multiple lavas containing Cl-excess. These areas are most apparent on the SMAR (Fig. 9a), where, within 10 km of a known hydrothermal site or plume, samples with Cl-excess are always seen, which are also some of the highest values observed along the ridge. High Cl-excess is also seen at the axial high of segment A2 (8°57′S) and the relatively shallow area in the north of A1 (7°51′S; Fig. 9a). Axial topographic highs and domes are volcanic structures with a higher average potential to host hydrothermal systems due to a higher magmatic activity, shown by detailed studies on the relation between hydrothermal occurrences and rift geomorphology (Fouquet 1997; Anderson et al. 2017). The axial high of segment A2 was earlier interpreted to be a remnant of past high magmatic activity (Devey et al. 2010) and there are currently no signs of hydrothermal activity (Devey et al. 2010; Haase et al. 2016). Therefore, we take the high Cl-excess in basalts at the axial highs of A2 and in the north of A1 to indicate that alteration of the crust took place there previously by ancient, now inactive hydrothermal systems, which makes these areas good targets for the search for extinct hydrothermal deposits.
Fig. 9

Locations of high Cl-excess samples (diamonds) and samples without Cl-excess (circles) and hydrothermal sites at the a SMAR and b Gakkel Ridge. Note in the vicinity of a known hydrothermal site there are always samples taken with elevated Cl-excess (> 25 ppm) with the exception of 85°E at Gakkel Ridge, where independent evidence suggests that the hydrothermal activity is probably younger than the magmas sampled (for discussion see text). The axial high at SMAR segment A2 (8°57′S) yielded Cl-excess samples but lacks any recent hydrothermal activity, which suggests that the high Cl-excess results from former, now inactive hydrothermal activity

At Gakkel Ridge, samples collected within 10 km of the Aurora hydrothermal field at 6°20′W and plume sites at 1°45′W and 37°E display high Cl-excess (> 25 ppm; Fig. 9b) consistent with the observations at the SMAR. Plume sites at 2°10′E, 7°30′E and 43°10′E display samples with high Cl-excess within the larger vicinity of 25 km, which is likely related to the highly uncertain locations for the origin of most of the hydrothermal plumes (Baker et al. 2004). The only exception among our samples are basalts that do not exhibit Cl-excess collected < 10 km from the present-day hydrothermal active site at 85°E (Fig. 8b; Baker et al. 2004; Pontbriand et al. 2012). But in detail, these sampled glasses come from visibly older lava flows than the young glassy basalts, believed to have erupted in 1999 that the present hydrothermal venting was linked to (Pontbriand et al. 2012). We therefore suggest that the 85°E system is a young hydrothermal system, that was established after the eruption of the here measured samples.

In summary, Cl-excess in erupted lavas always indicate that they have interacted with altered crust implying active or former hydrothermal circulation in the vicinity; the large amount of samples with Cl-excess show that hydrothermal alteration is widespread. Basalts with the highest Cl-excess (> 25 ppm) are related to places that display active or fossil hydrothermal venting in the vicinity and as such represents a guide that can aid to identify MOR segments with hydrothermal deposits.

Differences in Cl contamination between slow- and fast-spreading ridges

Our data show that Cl-contamination by assimilation of hydrothermally altered lithosphere is not restricted to fast-spreading ridges or ridges with a thick crust, contrary to the conclusions of Michael and Cornell (1998), but also occurs in several ultraslow- and slow-spreading ridges that show a range of morphologies and thicknesses. Nevertheless, differences in the amounts of Cl-contamination between fast- and slow-spreading ridges are clearly evident. Following Stroncik and Niedermann (2016) we assembled Cl and Cl/K data from the PetDB database (Lehnert et al. 2000) and used them to assess degrees of Cl-excess globally. Absolute MORB Cl-concentrations at slow-spreading ridges are overall considerably lower (only 14% has > 200 ppm Cl) than at fast- or medium-spreading ridges (36 and 31% has > 200 ppm Cl; Fig. 10a). Using an upper value of 0.09 for mantle Cl/K implies that at least 25% of the samples at slow-spreading ridges show Cl-excess (Fig. 10b), considerably less than observed at fast- and medium-spreading ridges (69 and 79%). The maximum Cl-excess on slow-spreading ridges in the PetDB database overlaps with our study (< 550 ppm, but mostly < 250 ppm) and is clearly lower than at fast- and medium-spreading ridges (mostly < 1000 ppm). The fact that only ca. 25% of slow-spreading magmas in PetDB are classified as showing Cl-excess compared to our 75% estimate for Gakkel and SMAR shows that using a global mantle Cl/K of 0.09 underestimates the samples with Cl-excess at slow-spreading ridges and emphasizes the need for both high-precision Cl determinations and good estimates of local mantle Cl/K values in these inherently Cl-poor magmas. Cl-excess in the PetDB data is not restricted to specific sections of slow-spreading ridges, e.g. with thick crust as suggested by Michael and Cornell (1998), as for example can be seen from the well-studied slow-spreading Northern Mid-Atlantic Ridge (Fig. 10c).
Fig. 10

Chlorine concentrations at different ridge types. a Frequency plots of Cl-contents and b Cl/K for slow-, medium- and fast-spreading ridges. Slow-spreading ridges show on average lower Cl-contents and Cl/K than medium- and fast-spreading ridges, nevertheless 23.9% of the slow-spreading MORB show Cl/K above a upper mantle limit value of 0.09 (dashed line). c Comparing samples with definite Cl-excess with hydrothermal sites on the most studied, slow-spreading Norther Mid-Atlantic Ridge (20°–65°N) shows that Cl-excess is present along the whole ridge and that where data are present hydrothermal sites are mostly associated with Cl-excess samples close by

Hence, Cl addition by hydrothermal contamination is present on all types of ridges, but shows a difference in the amount of Cl-excess in basalts from the two end-member type of ridges. While Cl-excess at fast-spreading ridges is high and present in most samples (e.g. Michael and Cornell 1998; Saal et al. 2002; le Roux et al. 2006; Kendrick et al. 2013), Cl-excess at slow-spreading ridges is more variable and significantly lower (Fig. 5a; Michael and Cornell 1998; van der Zwan et al. 2015; this paper). The lower absolute intensity of Cl-contamination in slow-spreading magmas implies either a lower degree of assimilation and/or a lower degree of lithospheric alteration. The less differentiated nature of magmas from slow- compared to fast-spreading ridges (higher average Mg#; Rubin and Sinton 2007) suggest that the magmas at slow-spreading ridges on average underwent less fractional crystallization and therefore have lost less heat during passage through the lithosphere. Whether that results in less extensive hydrothermal circulation and alteration or less assimilation of country rock, the net effect will be a lower average Cl addition into the slow-spread magmas. A decrease of hydrothermal alteration with depth (e.g. Dick et al. 2000; Coogan 2003) may also play a role in the lower Cl-excesses at slow-spreading ridges. The lower crustal crystallization depths shown here and deep assimilation of less altered rock will result in comparatively minor Cl addition compared to shallower magma chambers with chronic hydrothermal systems at fast-spreading ridges (< 3 km; le Roux et al. 2006; Kendrick et al. 2013) that assimilated more strongly altered crust. If less cooled by hydrothermal circulation, this deeper lithosphere needs to conductively loose its heat, consistent with thermodynamic models and inferred temperature gradients that indicate more effective conductive cooling at slow-spreading ridges (Reid and Jackson 1981; Niu and Hekinian 1997; Montési and Behn 2007). Slow-spreading ridges display a much higher compositional variety with heterogeneous amounts of alteration (e.g. Dick et al. 2000) and magma bodies are probably small and ephemeral, which implies that both assimilation and alteration processes are more localized, explaining that not all assimilated rock is altered at slow-spreading ridges leading to more variable amounts of Cl-excess in the samples. The large permanent magma lenses beneath fast-spreading ridges provide the opportunity for more extensive mixing of melts and assimilation of the crust over a larger scale, leading to a common chemical signature of assimilation of hydrothermally altered crust (cf. Saal et al. 2002; Kendrick et al. 2013).

Global implications of hydrothermal Cl contamination

The depth of hydrothermal circulation is important for the cooling of the lithosphere. With our data, we can demonstrate that there is a cooling mechanism by circulation of hydrothermal fluids down to the lower crust that can facilitate deep crystallization at slow-spreading ridges. This circulation most likely happens at all low spreading rates. This means crust is not merely formed at shallow levels but that lower crustal and lithospheric mantle magma chamber processes may be common at slow-spreading ridges and that oceanic crust can form in situ at all depths (cf. Kelemen et al. 1997; Kelemen and Aharonov 1998; Lissenberg et al. 2004; Grimes et al. 2008).

Hydrothermal circulation throughout the newly formed lithosphere can increase the Cl budget and to a smaller extent the H2O budget of an altered oceanic slab and thus the potential of Cl (and H2O) to subduct, although we cannot determine with the present data the Cl content and distribution in the altered plate and so cannot give a realistic estimate of its total Cl content. This has important consequences for global water and halogen budgets (see e.g. Kendrick et al. 2017). Although most other elements occur in too low concentrations in hydrothermal fluids to significantly change the composition of the deeper lithosphere, H- and S-isotopes, B and noble gasses may be significantly changed (Kent et al. 1999; Gillis et al. 2003; Kendrick et al. 2013; Labidi et al. 2014; Stroncik and Niedermann 2016). Any study of those elements and isotopes should take into account the effect of hydrothermal circulation and assimilation of hydrothermally altered rocks. In addition, melts that show Cl-excess at slow-spreading ridges likely interacted with the lithosphere and care should be taken when using their major element data for e.g. pressure or melt degree calculations.


Novel high-precision Cl measurements of basaltic glasses and melt inclusions in Ol and Pl from the slow- and ultraslow-spreading Southern Mid-Atlantic Ridge and Gakkel Ridge show up to 75% of these samples to have been affected by Cl-contamination, raising their Cl/Nb and Cl/K ratios above mantle values, which cannot be explained by magmatic processes such as degassing, melting, fractionation or source variations, or by seafloor processes. This contamination most likely occurs by the addition of hydrothermal brines to the magmas during the assimilation of hydrothermally altered country rock. Cl-contamination is shown to be a common process at (ultra-) slow-spreading ridges, in contrast to conclusions drawn from early studies based on lower precision Cl data. The compositional effects of crustal assimilation lead to disequilibrium and phase-equilibria-based estimates of magma fractionation depths, such as those provided by COMAGMAT, to be erroneous. Pressures calculated from Cpx–melt pairs are more reliable as equilibrium can be assessed and yield Cl-contamination depths (and thus minimum depths to which hydrothermal circulation penetrates) of at least 10 km for SMAR segment A2. This is near the base of the lower crust in this studied region. Melt inclusions with Cl-excess show assimilation of hydrothermal crust in magma chambers before crystallization at depth in magma reservoirs. Deep cooling by hydrothermal circulation and crustal assimilation is consistent with independent earlier evidence for lower crustal crystallization and melt–rock reactions and for extensive and deep faulting at slow-spreading ridges.

All MORB basalts that display Cl-excess were at some point in time affected by hydrothermal fluids, while the highest observed values here are associated with hydrothermal vent fields. Although globally both fast- and slow-spreading ridges display a similar proportion of samples with Cl-excess, slow-spreading ridges show both lower absolute Cl concentrations and maximum degrees of Cl-excess. This may result from either less alteration of the crust or less interaction of the magmas with the crust at slow-spreading ridges. The generally less fractionated nature of slow-spreading magmas provides support for the latter explanation—the former explanation would require a larger proportion of conductive cooling of the lower crust as spreading rates decrease. The detection of the effects of deep-seated hydrothermal circulation in slow-spreading crust has implications for the volatile and halogen budget of the oceanic plates and this needs to be taken into account when assessing the Cl, H2O, H-, S-, B- and noble gas isotopic signatures of global geochemical cycles.



We are very grateful to Mario Thöner for the extensive technical assistance at the EMP and to Dagmar Rau for the technical assistance at the LA-ICP-MS. Further, we like to thank Jan Fietzke (all GEOMAR) for the help with the modification of the Cl measurement method for the melt inclusions. The suggestion of three anonymous reviewers and editorial handling by Jochen Hoefs was greatly appreciated. We acknowledge generous financial support from the Jeddah Transect Project between King Abdulaziz University and Helmholtz-Center for Ocean Research GEOMAR that was funded by King Abdulaziz University (KAU) Jeddah, Saudi Arabia, under Grant no. (T-065/430).

Supplementary material

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© Springer-Verlag GmbH Germany 2017

Authors and Affiliations

  1. 1.Geomar Helmholtz Centre for Ocean Research KielKielGermany
  2. 2.Leibniz Universität Hannover, Institute of MineralogyHannoverGermany
  3. 3.GeoZentrum Nordbayern, Universität Erlangen-NürnbergErlangenGermany
  4. 4.Faculty of Marine ScienceKing Abdulaziz UniversityJeddahSaudi Arabia
  5. 5.Department of Earth and Atmospheric SciencesUniversity of HoustonHoustonUSA
  6. 6.Institute für Geowissenschaften, Christian-Albrechts-Universität KielKielGermany

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